New Mexico Bureau of Mines & Mineral Resources
A DIVISION OF

NEW MEXICO INSTITUTE OF MINING & TECHNOLOGY

Field guide to geologic excursions
in New Mexico and adjacent areasof Texas and Colorado
Prepared for the Geological Society of America
Rocky Mountain and South-Central Sections Annual Meeting,
Albuquerque, New Mexico, 21-24 April 1991

Edited by Betsy Julian' and Jiri Zidek2'Department of Geological Sciences, University of Texas at El Paso, El Paso, Texas 79968;'New Mexico Bureau of Mines & Mineral Resources, Socorro, New Mexico 87801

SOCORRO 1991

NEW MEXICO INSTITUTE OF MINING & TECHNOLOGY
Laurence H. Lattman, President
NEW MEXICO BUREAU OF MINES & MINERAL RESOURCES
Frank E. Kottlowski, Director and State Geologist
James M. Robertson, Associate Director
BOARD OF REGENTS

Ex Officio
Bruce King, Governor of New Mexico
Alan Morgan, Superintendent of Public Instruction
Appointed
Steve Torres, President, 1967-1997, AlbuquerqueCarol A. Rymer, M.D., President-Designate, 1989-1995, Albuquerque
Lt. Gen. Leo Marquez, Secretary/Treasurer, 1989-1995, Albuquerque
Robert 0. Anderson, 1987-1993, RoswellCharles Zimmerly, 1991-1997, Socorro
BUREAU STAFF
Full Time
ORIN J. ANDERSON, Geologist
RUBEN ARCHULETA, Metallurgical Lab. Tech.
AUGUSTUS K. ARMSTRONG, USGS Geologist
GEORGE S. AUSTIN, Senior Industrial Minerals Geologist
AL BACA, Maintenance Carpenter II
JAMES M. BARKER, Industrial Minerals Geologist
MARGARET W. BARROLL, Post-Doctoral Fellow
PAUL W. BAUER, Field Economic Geologist
ROBERT A. BIEBERMAN, Emeritus Sr. Petroleum Geologist
LYNN A. BRANDVOLD, Senior Chemist
RON BROADHEAD, Petrol. Geologist, Head, Petroleum Section
MONTE M. BROWN, Cartographic Drafter II
KATHRYN E. CAMPBELL, Cartographic Drafter II
STEVEN M. LATHER, Field Economic Geologist
RICHARD CHAMBERLIN, Economic Geologist
CHARLES E. CHAPIN, Senior Geologist
RICHARD R. CHAVEZ, Assistant Head, Petroleum Section
RUBEN A. CRESPIN, Garage Supervisor

NORMA J. MEEKS, Senior Pub./Bus. Office Clerk
BARBARA R. POPP, Chemical Lab. Tech. 11
MARSHALL A. REITER, Senior Geophysicist
JACQUES R. RENAULT, Senior Geologist
JAMES M. ROBERTSON, Senior Economic Geologist
JANETTE THOMAS, Cartographic Drafter II
SAMUEL THOMPSON III, Senior Petrol. Geologist
REBECCA J. TITUS, Cartographic Supervisor

Lois M. DEVLIN, Director, Bus./Pub. Office
ROBERT W. EVELETH, Senior Mining Engineer
Lois GOLLMER, Staff Secretary
IBRAHIM GUNDILER, Metallurgist
WILLIAM C. HANEBERG, Engineering Geologist
JOHN W. HAWLEY, Senior Env. Geologist
CAROL A. HJELLMING, Assistant Editor
GRETCHEN K. HOFFMAN, Coal Geologist
GLEN JONES, Computer Scientist/Geologist
ANN LANNING, Administrative Secretary
ANNABELLE LOPEZ, Petroleum Records Clerk
THERESA L. LOPEZ, Receptionist/Staff Secretary
DAVID W. LOVE, Environmental Geologist
JANE A. CALVERT LOVE, Editor
WILLIAM MCINTOSH, Research Geologist
CHRISTOPHER G. MCKEE, X-ray Facility Manager
VIRGINIA MCLEMORE, Geologist
LYNNE McNeil., Technical Secretary

JUDY M. VAIZA, Executive Secretary
MANUEL J. VASQUEZ, Mechanic I
JEANNE M. VERPLOEGH, Chemical Lab. Tech. II
ROBERT H. WEBER, Emeritus Senior Geologist
SUSAN J. WELCH, Assistant Editor
NEIL H. WHITEHEAD, III, Petroleum Geologist
MARC L. WILSON, Mineralogist
DONALD WOLBERG, Vertebrate Paleontologist
MICHAEL W. WOOLDRIDGE, Scientific Illustrator
JIRI ZIDEK, Chief Editor—Geologist

Research Associates
CHRISTINA L. BALK, NMT
WILLIAM L. CHENOWETH, Grand function, CO
PAIGE W. CHRISTIANSEN, Kitty Hawk, NC
RUSSELL E. CLEMONS, NMSUWILLIAM A. COBBAN, USGSAUREAL T. CROSS, Mich. St. Univ.
MARIAN GALUSHA, Amer. Mus. Nat. Hist.
LELAND H. GILE, Las Cruces

JEFFREY A. GRAMBLING, UNM
JOSEPH HARTMAN, Univ. Minn.
DONALD E. HATITN, Ind. Univ.
ALONZO D. JACKA, Texas Tech. Univ.
DAVID B. JOHNSON, NMT
WILLIAM E. KING, NMSUDAVID V. LEMONE, UTEPHOWARD B. NICKELSON, Carlsbad
LLOYD C. PRAY, Univ. Wisc.
ALLAN R. SANFORD, NMT
JOHN H. SCHILLING, Reno, NV
WILLIAM R. SEAGER, NMSURICHARD H. TEDFORD, Amer. Mus. Nat. Hist.
JORGE C. TOVAR R., Petroleos Mexicans

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WILLIAM BECKJENNIFER BORYTAROBERT KINGERNEST SCHARKAN, JR.
DAVID SIVILSPlus about 30 undergraduate assistants

Original Printing

Published by Authority of State of New Mexico, NMSA 1953 Sec. 63-1-4
Printed by University of New Mexico Printing Services, April 1991
Available from New Mexico Bureau of Mines & Mineral Resources, Socorro, NM 87801 Publishedas public domain, therefore reproducible without permission. Source credit requested.

Preface
This volume contains 10 field guides to trips scheduled to take place
before, during, and after the joint annual meeting of the GSA Rocky
Mountain and South-Central Sections in Albuquerque, New Mexico, on
21-24 April 1991. Most of the field guides combine road logs andtechnical papers, but Geomorphic and neotectonic evolution along themargin of the Colorado Plateau and Rio Grande rift, northern New
Mexico, by Gonzalez and Dethier, is a technical paper with field-trip stops
and topics of discussion only briefly listed in an appendix. The detailed
road log to this paper is Open file Report no. 374, available from the New
Mexico Bureau of Mines & Mineral Resources, Socorro, New Mexico
87801, tel. (505) 835-5147.
The field guides are arranged in the order in which the trips were announced (GSA News & Information, v. 12, no. 10, October 1990); however,three of those announced are not included. One is Educational geology of the
Albuquerque area, a pre-meeting, one-day trip open only to schoolteachers,
for which no field guide has been submitted. The other two are post-meeting
trips, Geology of the Valles caldera and Jemez volcanic field (one day) andCenozoic magmatism and tectonics of the southeastern Colorado Plateau,
New Mexico (three days), for which field guides have already appeared in the
New Mexico Bureau of Mines & Mineral Resources Memoir 46 (1989), andthe field-trip participants thus can be supplied with reprints from thatpublication.
Fifty-five authors have contributed to the volume. Because of the large
number of authors, their geographic separation, and the short time from
submittal to publication (two field guides were made available to us onlyin mid-February!), it has not been possible to completely standardize theformat and to improve on the quality of some of the figures and tables.
However, we have done everything possible to turn out a good-quality
publication, and we hope our effort will be viewed with the understandingthat we had to work fast.
We would like to take this opportunity to thank the authors for theircontributions and the University of New Mexico Printing Services for an
outstanding performance in all phases of production of this volume.
Betsy Julian
Jiri Zidek

New Mexico Bureau of Mines & Mineral Resources Bulletin 137, April 1991

1

Examples of modern rift volcanism and Proterozoic anorogenic
magmatism: The Potrillo volcanic field of southern New Mexico andthe Franklin Mountains of west Texas
E. Y. Anthony', C. G. Barnes', W. Chen', J. M. Hoffer', G. R. Keller', K. M. Marsaglia¹,
V. T. McLemore¼, J. M. Seeley', M. A. Seward', W. M. Shannon', and W. F. Thomann''Department of Geological Sciences, University of Texas at El Paso, El Paso, Texas 79968; ²Department of Geology, Texas Tech University, Lubbock, Texas 79409;³
Department of Geosciences, Texas A&I University, Kingsville, Texas 78363; 4New Mexico Bureau of Mines & Mineral Resources, Socorro, New Mexico 87801

Introduction
This article describes a three-day field trip which origi- regional Proterozoic geology. The second two days willnates in El Paso, Texas, and terminates in Albuquerque, focus on features of the Rio Grande rift. We will visit the
New Mexico. The first day we will present an overview of Quaternary Potrillo volcanic field, with stops at a classic
the present-day Rio Grande rift and also examine the Pro- shield volcano — the Aden Crater — and two maars,
terozoic sedimentary section in the Franklin Mountains, including Kilbourne Hole. The second night will be spent
the 1.1 Ga Red Bluff Granite which intrudes it, and coeval in Las Cruces. The third day we will travel totrachyte lavas and rhyolitic ignimbrites in the Thunderbird Albuquerque, stopping to examine the tectonic style andGroup. The purpose of these stops is to provoke discussion geologic features along the way.
of tectonic models for this sequence within the context ofDay 1: The Proterozoic section in the Franklin Mountains
C. G. Barnes, K. M. Marsaglia, J. M. Seeley, M. A. Seward, W. M. Shannon, and W. F. Thomann
The focus of Day 1 is an overview of Proterozoic sections
exposed in the Franklin Mountains, Texas (Figs. 1, 2), and ageneral introduction to the Rio Grande rift (see IGC guidebook for Day 3, Callender et al., 1989). To access this region,
take the northbound US-54 exit from I-10. Note that US-54
becomes FM 3255 (the North—South Freeway). From US54 (FM 3255) take the Loop 375 (Woodrow Bean TransMountain Road) exit, and follow road log below. Please use
caution when examining the unstable roadcuts and outcrops
along Loop 375.
Mileage

0.0 Start at intersection of Loop 375 (Woodrow Bean
Trans-Mountain Road) and FM 3255 (North—South
Freeway). Proceed west on Loop 375 past Wilderness Park Museum to Stop 1.
0.9 Stop 1: Roadside next to Indian Spring Canyon. To
the east lies the Hueco bolson, an extensional basin
filled with over 3 km of Cenozoic sediments (seeStop 3, IGC guidebook). The low knob to the northis a relatively fresh exposure of the main phase of the
Red Bluff Granite (Nelson, 1940), which is a coarsegrained biotite—amphibole granite. Hike due northfor 0.5 mi, then proceed west, up dirt road 0.5 miinto Indian Spring Canyon. The entrance of Indian
Spring Canyon is underlain by biotitearfvedsonite
granite. It occurs as a sill hosted by the Red Bluff
Granite and crops out continuously for
approximately 1 mi. The biotite—arfvedsonite
granite is locally banded or layered (i.e. line rock).
Apophyses from this sill extend southeast to Trans-

Mountain Road and northwest to an abandoned tinmine locality.
1.5 Stop 2: First roadcut on Loop 375 West. This outcrop consists of blocks of Castner Marble (see description below) and metabasalt (geochemicallydistinct from the Mundy Breccia; Thomann and
Hoffer, 1985) that are hosted and intensely injectedby a medium-grained to pegmatitic leucocratic biotite granite (Harbour, 1972). A block of the porphyritic quartz syenite (with Lanoria Quartzite still
attached) is exposed on the northeastern end of theoutcrop. The lack of internal stratigraphic ordersuggests that these blocks are stope blocks. In
addition, questionable sedimentary structures (mudcracks) within the xenolith of Castner Marble at thewest end of the roadcut suggest that this block may
be overturned. The leucocratic biotite granite isprobably a late phase of the Red Bluff Granite (seeShannon and Barnes, this volume). Dikes of biotite—arfvedsonite granite cut all other rock types.
2.5 Stop 3: Pull off into parking area just before third
roadcut on Loop 375 West. At this location, Proterozoic metasediments (Castner Marble and Lanoria Quartzite) and metavolcanics (Mundy Breccia
and diabase sills) comprise a roof pendant within the
Red Bluff Granite complex. The intrusive contact
between the Red Bluff Granite complex and thislarge xenolith is well-exposed on the slope to thenorth. The light-gray, bedded unit is the oldest
formation, the Castner Marble, which consists ofthinly laminated to thickly bedded, medium- tocoarse-grained marble with alternating layers ofHoffer, 1976; LeMone, 1982; LeMone and Cornell,
1988; Thomann and Hoffer, 1989). Along this roadcut the Castner Marble/Mundy Breccia contact ispoorly exposed, but outcrops just north of the roadcut show the disrupted nature of this contact (seeThomann, this volume, for further discussion of theMundy Breccia).
2.8 Stop 4: Fourth roadcut on Loop 375 West. The
granitic sill exposed at this location has been informally named porphyritic microgranite (McAnulty,1967). We suggest that this terminology should beabandoned (see Shannon and Barnes, this volume).
The sill was emplaced into the lower member ofthe Lanoria Quartzite and has chilled margins 1-5
m thick. The center of the sill is a porphyritic biotite—amphibole quartz syenite and is probably cumulate. The chilled margins are granitic porphyry.
Numerous Proterozoic basaltic dikes cut the sill.
3.1 Stop 5: West end of fourth roadcut on Loop 375 West.
Exposures of the middle Proterozoic Lanoria
Formation can be seen to the north. This formationconsists of approximately 750 m of quartzitic sandstones, siltstones, and mudstones which have been
intruded and contact-metamorphosed by granites of
the Red Bluff Granite complex to at least the albiteepidote-hornfels facies. Due to the low degree ofmetamorphism, sedimentary structures and textures
are relatively well preserved. Harbour (1960) divided the formation into three unnamed members
based primarily on topographic constraints (Fig. 2),
whereas Seeley (in preparation) subdivides thesethree members into six members (L1 through L6;
Fig. 2) based on lithology and depositional environment. According to Seeley (in preparation), theserocks were deposited in shallow, subtidal-shelf toestuarine environments (members L1, L3 and L5:
inner shelf; members L2 and L4: outer shelf; member L6: estuarine lagoon behind a protective tidal
shoal). The members of the Lanoria represent pulsesof transgression and regression across a gently subhornfels. Also exposed on this slope are diabase sills
within the Castner Marble, and the Mundy Breccia
which unconformably overlies the Castner Marble
and pinches out just below the ridge crest. The cliffforming unit above the Mundy Breccia is thelowermost member of the Lanoria Quartzite (LI , seeStop 5). In the valley to the right, algal stromatolitic
mounds and Conophyton-type stromatolites havebeen described in a lower block of the Castner
Marble (Toomey and Babcock, 1983). The Castner
Marble is exposed along the roadcut to the left of theparking area. Hoffer (1976) described two
metamorphic facies in the Castner Marble: (1)hornblende-hornfels facies and (2) K-feldspar—cordierite-hornfels facies, which places the conditionsof metamorphism at approximately 515°C and 1000
bars of pressure. A distinctive skarn zone in theCastner Marble is exposed at the eastern edge of thisroadcut. Further up section, an enigmatic breccia
pipe (or clastic dike?) cuts across the Castner
Marble. The amount of hornfels (meta-shale), softsediment deformation and edgewise, intraformational conglomerates increases towards the top of
the formation (Denison and Hetherington, 1969;

3
siding platform. Facies distributions and paleocurrent data suggest that the platform sloped northward
and that the shelf gradient was relatively low. At
this stop, the gray L2 member of the Lanoria Formation is exposed on the lower (gentle) slope, andthe overlying middle members (L3 through L5;
white quartzite beds and a gray, fine-grained unit)
form the cliff along the ridge.
4.3 Stop 6: Carefully cross over gravel median and enter
large parking lot on opposite side of road. Massive
outcrops of the Thunderbird Group (Thomann,
1980, 1981) can be seen up ahead in roadcuts along
Smugglers Pass. From parking area, proceed south
along marked hiking trail into canyon, where thethree formations of the Thunderbird Group are well
exposed. The Coronado Hills Conglomerate is thebasal unit that forms prominent, relatively steep
cliffs. Units of the Smugglers Pass Formation occur
above the basal conglomerate and include a lowermost melanocratic porphyritic trachyte that grades
into a lighter-colored (pink) porphyritic trachyte,followed by a relatively thick, light-colored tuffaceous sandstone that is also a cliff-former. A thinporphyritic trachyte overlies this sandstone. The
thickest and uppermost formation, the Tom Mays
Park Formation, is a recrystallized, metamorphosed
ignimbrite (originally a crystal[quartz, K-feldspar]vitric tuff). In outcrop and hand specimen, the Tom
Mays Park Formation superficially resembles a porphyritic rhyolite, but thin-section examination reveals its true pyroclastic nature. Despite contact
metamorphism from intrusions of the Red Bluff
Granite complex, all igneous and sedimentary textures and structures are well preserved at this locality and throughout the Franklin Mountains.
5.1 Stop 7: Carefully cross over gravel median andcontinue west on Loop 375 through Smugglers
Pass. Carefully cross back over gravel median andenter large parking lot on opposite side of road forscenic overview of Mesilla basin (see Stops 4 and
5, IGC guidebook).

References
Callender, J. F., Hawley, J. W., and Keller, G. R., 1989, Rio Grande rift:
28th International Geological Congress, Guidebook to Field Trip T318,
37 pp.
Denison, R. E., and Hetherington, E. A., 1969, Basement rocks in far west
Texas and south-central New Mexico: New Mexico Bureau of Mines &Mineral Resources, Circular 104: 1-16.
Harbour, R. L., 1972, Geology of the northern Franklin Mountains, Texas
and New Mexico: U.S. Geological Survey, Bulletin 1298: 129 pp.
Hoffer, R. L., 1976, Contact metamorphism of the Precambrian Castner
Marble, Franklin Mountains, El Paso County, Texas: Unpublished MS
thesis, University of Texas at El Paso, 77 pp.
LeMone, D. V., 1982, Stratigraphy of the Franklin Mountains, El Paso
County, Texas and Doña Ana County, New Mexico: El Paso GeologicalSociety, Symposium on the Franklin Mountains, pp. 42-72.
LeMone, D. V., 1989, Stratigraphy and structure of the Franklin Mountains, west Texas; in Structure and stratigraphy of Trans-Pecos Texas:
28th International Geological Congress, Field Trip Guidebook T317:
55-64.
LeMone, D. V., and Cornell, W. C., 1988, Franklin Mountains, Texas and
New Mexico: El Paso Geological Society and American Association ofPetroleum Geologists, Stratigraphy Field Guide, 43 pp.
McAnulty, W. N., Jr., 1967, Geology of the Fusselman Canyon area,
Franklin Mountains, El Paso County, Texas: Unpublished MS thesis,
University of Texas at Austin, 79 pp.
Nelson, L. A., 1940, Paleozoic stratigraphy of Franklin Mountains, west
Texas; in DeFord, R. K., and Lloyd, E. R. (eds.), West Texas—New
Mexico Symposium: American Association of Petroleum Geologists,
Bulletin, 24(1): 157-172.
Seeley, J. M. (in preparation), Stratigraphy, petrology, and depositional
environment of the Middle Proterozoic Lanoria Formation, Franklin
Mountains, El Paso County, Texas: Unpublished MS thesis, Texas A&l
University, Kingsville.
Thomann, W. F., 1980, Petrology and geochemistry of the PrecambrianThunderbird Formation, Franklin Mountains, El Paso County, Texas:
Unpublished PhD dissertation, University of Texas at El Paso, 162 pp.
Thomann, W. F., 1981, Ignimbrites, trachytes, and sedimentary rocks of
the Precambrian Thunderbird Group, Franklin Mountains, El Paso,
Texas: Geological Society of America, Bulletin, 92: 92-100.
Thomann, W. F., and Hoffer, J. M., 1985, Petrology and geochemistry of
the Precambrian Mundy Breccia, Franklin Mountains, Texas: Texas
Journal of Science, 36: 267-281.
Thomann, W. F., and Hoffer, R. L., 1989, Stratigraphy, sedimentary structures and petrology of the Middle Proterozoic Castner Marble, Franklin
Mountains, far west Texas: University of Wyoming, Contributions toGeology, 27: 33-39.
Toomey, D. F., and Babcock, J. A., 1983, Precambrian and Paleozoic
algal carbonates, west Texas and southern New Mexico: Colorado
School of Mines, Professional Contribution 11: 345 pp.

Geology of the middle Proterozoic Thunderbird Group, Franklin Mountains, El Paso, Texas
William F. Thomann

Introduction
The middle Proterozoic Thunderbird Group occurs in the
Franklin Mountains, north of El Paso, Texas. The Franklins
are a north—south-trending, fault-block range located in thesoutheastern part of the Basin and Range province. Westdipping middle Proterozoic through Paleozoic age strata are
well exposed along the steep, east-facing scarps of the range.
Continuous outcrops of the Thunderbird Group extend fromMundys Gap, located just east of Franklin Mountains State
Park (formerly Tom Mays Memorial Park), to South Franklin
Mountain. The Thunderbird Group is the youngest of astratigraphic succession of middle Proterozoic rocks. This
succession consists of, from oldest to youngest, the Castner
Marble (Denison and Hetherington, 1969; Hoffer, 1976;
Thomann and Hoffer, 1989), formerly called Castner Limestone (Harbour, 1960, 1972), the Mundy Breccia (Harbour,
1960, 1972; Thomann and Hoffer, 1985), the Lanoria
Quartzite (Richardson, 1909), and the Thunderbird Group(Thomann, 1980, 1981). The Red Bluff Granite complex
intruded all of the middle Proterozoic sedimentary andvolcanic rocks. Precambrian(?) diabase dikes subsequently
intruded all of these. Isotopic dating of the volcanics of theThunderbird Group, and of the granites, indicates an age of —950 Ma (Denison and Hetherington, 1969). Denison andHetherington (1969) stated that the granites and volcanics
were contemporaneous (comagmatic?) because they fit asingle isochron. More recent age determinations by Norman
et al. (1987) and Copeland and Bowring (1988) indicate thegranites are about 1.1 Ga.
Despite devitrification of all glass and subsequent contact

4
metamorphism, relict igneous and sedimentary textures arewell preserved in units of the Thunderbird Group. Thus, theThunderbird Group provides evidence on the origin of
these middle Proterozoic igneous and sedimentary rocks in
the Franklin Mountains.
The Thunderbird GroupThe Thunderbird Group is divided into three formations
based upon distinct lithologies and consists of the lower
Coronado Hills Conglomerate, the middle Smugglers Pass
Formation composed of volcanic and sedimentary rocks,
and the upper Tom Mays Park Formation composed almost
entirely of rhyolitic ignimbrites. The term "ignimbrite," asdefined by Sparks et al. (1973), refers to a pyroclastic rockoriginally composed of pumice and glass shards, regardlessof degree of welding. The type locality of the Thunderbird
Group is in the Franklin Mountains, —914 m northeast of
Smugglers Pass in Fusselman Canyon at latitude 31°53'06"
and longitude 106°29'34".
Coronado Hills Conglomerate
The Coronado Hills Conglomerate consists of well-rounded
cobbles and pebbles of quartzite, quartz—sericite siltstone,
shale, chert, jasper, ignimbrite, porphyritic trachyte, and rare
marble and granite in a recrystallized matrix composed ofsand-sized quartz, alkali-feldspar, devitrified pumice, and glass
shards. Most of the sedimentary clasts were presumably
derived from erosion of the Lanoria Quartzite and Castner
Marble, whereas the volcanic and granite clasts were derived
from erosion of older but unknown source rocks. Thickness ofthe conglomerate ranges from 11 to 27 m. Structures such as
horizontal bedding, scour fills, and cross-bedding are visible inoutcrops throughout this formation.
Smugglers Pass Formation
Melanocratic porphyritic trachyte member—A melanocratic porphyritic trachyte, up to 46 m thick, is the lowermost unit of the Smugglers Pass Formation. The dark-grayto olive-black groundmass is due to the presence ofabundant, very fine-grained amphibole. Some other trachytes, however, contain abundant, very fine-grained chlorite and opaque minerals which are indistinguishable fromamphibole except in thin-section. Phenocrysts are predominantly albitized microcline perthites with very minor
amounts of plagioclase; some perthites have albite rims.
Glomerophenocrysts of anhedral feldspars are common; afew possess albite rims. The few plagioclase phenocrysts arealtered entirely to an assemblage of albite + epidote ±
calcite. Hence, the original plagioclase composition is notknown. The albite + epidote ± calcite assemblage is found in
no other rock type throughout the Thunderbird Group.
Porphyritic trachyte member—The porphyritic trachyte
is up to 61 m thick. Two facies are present. The grayish-red
trachytes are similar to the melanocratic trachytes except for
the absence of very fine-grained amphibole in the
groundmass. The presence of abundant, fine-grained opaquesin some trachytes produces a dark-red groundmass. Albitized
microcline perthite is the predominant phenocryst and someof these grains occur as glomerophenocrysts of randomly
oriented feldspars. Albite rims some grains andglomerophenocrysts of microcline perthite. Quartz and albite phenocrysts are rare. Relict zoning is present in a few ofthe least altered (albitized) microcline-perthite phenocrysts.
Groundmass minerals are a very fine-grained, equi

granular assemblage of albite, K-feldspar, and some quartz.
Tuffaceous sandstone and conglomerate member—
The tuffaceous sandstones and conglomerates are up to 27
m thick. The tuffaceous sandstones and the matrix of the
tuffaceous conglomerates consist of detrital particles ofquartz, pumice, and lithic fragments, and minor amounts ofperthite, albitized perthite, and magnetite. Quartz grains are
angular to well rounded; feldspar, lithic, and pumice grains
are angular to subrounded. The framework of the tuffaceous
conglomerates is composed of rounded boulders, cobbles,and pebbles of trachyte, silicified ignimbrite, and minor
amounts of quartzite. Primary sedimentary structures visible
in outcrops include horizontal bedding and low-angle crossbedding.
Silicified ignimbrite member—The silicified ignimbrites
vary in thickness from one to about 6 m. Microclineperthitephenocrysts and rare quartz phenocrysts occur within a darkgray groundmass that is partly to completely replaced bysilica. In thin-section, compacted pumice and glass shards
draped partly around phenocrysts indicate that these tuffs aredensely welded. All original glass devitrified to a very finegrained assemblage of quartz and alkali-feldspar. Groundmass minerals consist of microphenocrysts of microcline
perthite and minor amounts of primary (volcanic) quartz. A
few ash-fall deposits occur throughout the silicified ignimbrites. These ash deposits are no more than several centimeters thick and are locally silicified.
Tom Mays Park Formation
Ignimbrites—The ignimbrites of the Tom Mays Park
Formation attain a maximum thickness of 168 m. An erosional surface at the top of the formation indicates that it
was originally thicker. Individual flow units have not beenrecognized due to the uniform texture of the groundmass
throughout the formation. The uniform texture wasproduced in part by devitrification of glass and in part bysubsequent contact metamorphism caused by intrusions of
the Red Bluff Granite complex.
In thin-section, textural alterations in the groundmass
vary from those ignimbrites which retained original
outlines of glass shards and pumice to those which containlittle or no evidence of shards and pumice. The least alteredignimbrites retained not only evidence of the pyroclastic
origin of these rocks but also evidence of dense welding,even though the glass recrystallized to a very fine-grained
assemblage of alkali-feldspar, quartz, and opaques. Most of
the ignimbrites in the formation contain little or noevidence of shards and pumice because of textural andmineralogical changes in the groundmass.
Pumice occurs as compacted lenses and irregularly shaped
shreds consisting of a fine-grained equigranular assemblageof quartz and alkali-feldspar. The coarser-grained texture andlighter color of the pumice relative to the groundmass permit
detection of pumice in even highly recrystallized ignimbrites.
The longest dimension of compacted pumice is less than 20
times the length of the smallest dimension measured
perpendicular to bedding. According to Ross and Smith(1961), this suggests flattening due to compaction and notelongation due to stretching.
Almost all ignimbrites are crystal-vitric tuffs. Phenocrysts
consist of quartz and alkali-feldspar, and some quartz phenocrysts exhibit resorption with either rounded edges and/ or
embayed features. Most feldspar phenocrysts are too alteredto identify the type of feldspar. In the least altered5
ignimbrites, however, the feldspar phenocrysts are microcline perthites. In most other ignimbrites, the feldspars arehighly albitized along rims, cracks, and cleavages. The
groundmass consists of a very fine-grained, equigranularassemblage of microcline perthite, quartz, and albite.
Porphyritic rhyolites—Porphyritic rhyolites occur as
widely scattered dikes which intrude the ignimbrites of the
Tom Mays Park Formation. In volume, the dikes are insignificant compared to the rhyolitic ignimbrites in this formation. There are no rhyolite lava flows recognizedanywhere within the Tom Mays Park Formation. The
porphyritic rhyolites contain quartz and alkali-feldspar
phenocrysts in a fine-grained groundmass of quartz,
microcline perthite, and albite.
Contact metamorphism
Elevated temperatures accompanying granite intrusionsof the Red Bluff Granite complex metamorphosed rocks of
the Thunderbird Group to the albite—epidote-hornfels
facies. Absence of penetrative deformation and thepreservation of igneous and sedimentary textures indicatethat lithostatic pressure was low. Evidence of metamorphic
grade is based on thin-section identification of diagnostic
metamorphic-mineral assemblages in particular rock typesof the Thunderbird. The melanocratic porphyritic trachyte
provides the most critical evidence of metamorphic grade.
This trachyte contains coexisting albite + epidote + chlorite+ sericite ± calcite. Coexisting metamorphic minerals in theother rock types of the Thunderbird Group are chlorite +albite + sericite, an assemblage which is not strictly
diagnostic of the albite—epidote-hornfels facies.
Conclusions
Recognition of ignimbrites in the middle Proterozoic rocksof the Franklin Mountains, Texas, confirms the suggestions
made by Sellards et al. (1932) and Harbour (1960, 1972)
concerning the possible pyroclastic origin of the volcanic
rocks generally referred to as "rhyolite porphyry." The ignimbrites of the Tom Mays Park Formation in the Thunderbird Group resemble, in outcrop and hand specimen, arhyolite porphyry. Detailed petrographic analysis is required
to recognize the pyroclastic origin.
The Franklin Mountains provide a unique setting in the
study of the middle Proterozoic rocks that compose theThunderbird Group. The formations of the Thunderbird
Group are the only late middle Proterozoic metavolcanicand volcaniclastic rocks in the west Texas area which (1)are well exposed in a continuous sequence, (2) are virtually
undeformed, and (3) possess well preserved volcanic andsedimentary textures.
The nearest outcrops of rhyolitic volcanics of age andlithology similar to the Thunderbird Group are located at
Pump Station Hills, Hudspeth County, Texas, —70 mi (113km) east of El Paso, Texas. Rb/Sr ages based on feldspar
and whole-rock analyses for a "rhyolite porphyry" are 996
Ma and 959 Ma, respectively (Wasserburg et al., 1962).
U—Pb ages of zircons extracted from a "rhyolite porphyry"
are older, ranging from 1005 to 1175 Ma (Wasserburg etal., 1962). Thin-section examination of the "rhyolite" and"rhyolite porphyry" reveals mineralogy and texture almost
identical to the altered ignimbrites of the Tom Mays Park
Formation in the Thunderbird Group. Thus, the volcanics at
the Pump Station Hills appear to be correlative to thevolcanics in the Franklin Mountains.
References
Copeland, P., and Bowring, S., 1988, U—Pb zircon and 40Ar/39Ar ages
from Proterozoic rocks, west Texas (abs.): Geological Society of America, Abstracts with Programs, 20: 95-96.
Denison, R. E., and Hetherington, E. A., Jr., 1969, Basement rocks in far
west Texas and south-central New Mexico: New Mexico Bureau of
Mines & Mineral Resources, Circular 104: 1-16.
Harbour, R. L., 1960, Precambrian rocks at North Franklin Mountain,
Texas: American Association of Petroleum Geologists, Bulletin, 44:
1785-1792.
Harbour, R. L., 1972, Geology of the Northern Franklin Mountains, Texas
and New Mexico: U.S. Geological Survey, Bulletin 1298, 129 pp.
Hoffer, R. L., 1976, Contact metamorphism of the Precambrian Castner
Marble, Franklin Mountains, El Paso County, Texas: Unpublished MS
thesis, University of Texas at El Paso, 77 pp.
Muehlberger, W. R., Denison, R. E., and Lidiak, E. G., 1967, Basement
rocks in continental interior of United States: American Association ofPetroleum Geologists, Bulletin, 51: 2351-2380.
Norman, D. I., Condie, K. C., Smithy, R. W., and Thomann, W. F., 1987,
Geochemical and Sr and Nd isotopic constraints on the origin of the late
Proterozoic volcanics and associated tin-bearing granites from theFranklin Mountains, west Texas: Canadian Journal of Earth Sciences,24: 830-839.
Richardson, G. B., 1909, Descriptions of the El Paso quadrangle, Texas:
U.S. Geological Survey, Atlas, Folio 166, 11 pp.
Ross, C. S., and Smith, R. L., 1961, Ash flow tuffs: Their origin, geologic
relations, and identification: U.S. Geological Survey, Professional Paper
366, 81 pp.
Sellards, E. H., Adkins, W. S., and Plummer, F. B., 1932, The geology ofTexas: University of Texas Bureau of Economic Geology Bulletin,
Publication 3232: 1-44.
Sparks, R. S. J., Self, S., and Walker, G. P. L., 1973, Products ofignimbrite eruptions: Geology, 1: 115-118.
Thomann, W. F., 1980, Petrology and geochemistry of the PrecambrianThunderbird Formation, Franklin Mountains, El Paso County, Texas:
Unpublished PhD dissertation, University of Texas at El Paso, 162 pp.
Thomann, W. F., 1981, Ignimbrites, trachytes, and sedimentary rocks of
the Precambrian Thunderbird Group, Franklin Mountains, El Paso,
Texas: Geological Society of America, Bulletin, 92: 94-100.
Thomann, W. F., and Hoffer, J. M., 1985, Petrology and geochemistry of
the Precambrian Mundy Breccia, Franklin Mountains, Texas: Texas
Journal of Science, 36: 267-281.
Thomann, W. F., and Hoffer, R. L., 1989, Stratigraphy, sedimentary structures, and petrology of the Middle Proterozoic Castner Marble, Franklin
Mountains, far West Texas: University of Wyoming, Contributions toGeology, 27: 33-39.
Wasserburg, G., Wetherill, W. W., Silver, L. T., and Flawn, P. T., 1962, A
study of the ages of the Precambrian of Texas: Journal of GeophysicalResearch, 67(10): 4021-4047.

Petrology of the middle Proterozoic Mundy Breccia, Franklin Mountains, Texas
William F. Thomann
Introduction
The Mundy Breccia is one of four Proterozoic stratigraphic units exposed in the Franklin Mountains. These
include, from oldest to youngest, the Castner Marble (Hoffer, 1976; Thomann and Hoffer, 1989), Mundy Breccia(Harbour, 1960, 1972; Thomann and Hoffer, 1985), Lanoria
Quartzite (Richardson, 1909), and Thunderbird Group(Thomann, 1980, 1981). The middle Proterozoic Mundy

6
Breccia crops out on the east side of the Franklin Mountains
at four isolated outcrops; the largest one is located adjacent
to and north of the Trans-Mountain Road. The Mundy Breccia has not been studied in any detail and previous workers
described only briefly its petrography, but they suggested
possible origins: volcanic agglomerate (Harbour, 1960,
1972), weathered basalt lava flow or sill (Harbour, 1960,
1972), basalt flow breccia (McAnulty, 1967), and surface orsubmarine flow breccia (Denison and Hetherington, 1969).
Contact metamorphism of all Proterozoic sedimentary andvolcanic rocks in the Franklin Mountains resulted in mineralogical alteration with virtually no change in original
textures and structures. Heat derived from intrusive granitesof the Red Bluff Granite complex (Ray, 1982; McCutcheon,
1982) metamorphosed all Precambrian stratigraphic units,with the highest-grade metamorphic rocks occurring at andnear the intrusive margins and the lowest-grade metamorphic
rocks located well away from the granite contacts (Harbour,
1960, 1972; McAnulty, 1967; Dye, 1970; Hoffer, 1976;
Denison and Hetherington, 1969; Thomann, 1980, 1981;
Thomann and Hoffer, 1989).
Stratigraphy
The Mundy Breccia varies in thickness from zero to 64 m
at the largest outcrop north of the Trans-Mountain Road andtrends approximately north—south with an average dip of
27° west. The breccia lies unconformably on the Castner
Marble along a contact marked by a 3 to 20 cm skarn zone.
This zone formed as a result of thermal metamorphism from
intrusive granites. Chemical reactions that occurred between
the basalt and marble resulted in the formation of an equilibrium mineral assemblage of primarily calc-silicates (Hoffer, 1976).
The breccia filled-in channels and also produced depressions by differential loading as it flowed onto the semiconsolidated carbonate sediments. Evidence of disrupted beddingin the upper portion of the Castner Marble includes load casts,
intraformational breccias, imbricate structures, and folds.
Along strike, the upper contact of the Mundy Breccia with the
Lanoria Quartzite is sharp and uniform. The upper contact isan erosional surface and, therefore, the original thickness ofthe breccia is not known.
The framework of the Mundy Breccia consists almost
entirely of angular to subrounded basalt clasts ranging insize from granule to boulder. The fabric varies considerably,both vertically and laterally, from a clast-supported framework with little matrix to a very open framework with clasts
supported by matrix. There is no evidence of superposed
mudflows.
The breccia is unsorted and unstratified. Spheroidally
weathered clasts in various stages of decomposition constitute the matrix of the upper 4.5 m of the breccia. These
weathered clasts provide evidence of weathering, and probable erosion, prior to burial by overlying sediments of theLanoria Formation. The intensity of weathering diminishes
gradually beneath the quartzite—breccia contact; the lower
basalt fragments retain their original angular to subroundedshapes.
Denison and Hetherington (1969) interpreted the Mundy
Breccia as a mudflow breccia but they, as well as otherauthors, could not determine whether the breccia was asurface or submarine flow. The Mundy Breccia has characteristics of a mudflow breccia that includes granule- toboulder-sized fragments derived from weathering and ero

sion of a basalt. More recently, Thomann and Hoffer(1985) concluded that the Mundy Breccia is an epiclastic
volcanic breccia.
Petrography
Basalt clasts in the Mundy Breccia are predominantly
aphanitic; a few are amygdaloidal. Thin-sections of theclasts reveal a metamorphic mineralogy in the basalts withwell preserved relict intersertal (glassy) textures seen in
only a few samples. The very fine-grained nature of muchof the groundmass in the clasts and matrix prevents an
accurate point-count determination of percentages ofmetamorphic minerals.
Cobble- and boulder-size clasts exhibit the best preserved
relict intersertal (and/or amygdaloidal) textures. In thesebasalts, plagioclase crystals retain their original lath shapes,
although they altered to albite and/or sericite. Microphenocrysts now composed of fine-grained actinolite, minor biotite,plagioclase, and less commonly sericite, are pseudomorphic
after olivine and pyroxene. Former glass (tachylite?) altered toa very fine-grained aggregate of feldspar (albite?), sericite,
magnetite, tremolite, actinolite, biotite, and lesser amounts of
epidote, clinozoisite, and chlorite. In some samples, finegrained magnetite is so abundant that extensive portions of athin-section are opaque to transmitted light. Those basalts withthe rarer amygdaloidal texture have vesicles filled with veryfine-grained muscovite, biotite, and minor amounts of coarsergrained quartz and feldspar. The well preserved textures andthe relatively low concentrations of olivine and pyroxene (asshown by pseudomorphs) suggest that these rocks aretholeiitic olivine basalts. The remaining basalts show a wide
range of textural variations from those with some traces offormer relict igneous textures (i.e. lath shapes of former
plagioclase) to those so thoroughly recrystallized that noevidence of former textures remains. The relative amounts ofmetamorphic minerals vary considerably from one sample tothe next throughout the breccia. Generally, in the intensely
recrystallized and usually coarser-grained metabasalts, thepredominant replacement minerals are actinolite, biotite,
epidote, and/or clinozoisite.
Matrix minerals in the Mundy Breccia include actinolite,
tremolite, biotite, epidote, magnetite, albite, chlorite, clinozoisite, muscovite, quartz, microcline, and some zoned
tourmaline. Small patches, streaks, lenses, or fracture fillings of actinolite, biotite, epidote, quartz, and magnetite arecoarse enough to identify in some hand specimens by the
unaided eye or hand lens. Locally, some portions of thematrix are composed predominantly of quartz with muscovite, biotite, albite, actinolite, microcline, and magnetite
in subordinate amounts. However, there is no uniformity inthe distribution of mineral assemblages found throughout
the matrix in the Mundy Breccia.
The mineralogy and texture of the spheroidally weatheredbasalt in the upper 4.5 m of the breccia are similar to theunderlying metabasalts. Although original minerals are nolonger present, relict textures of the weathered rinds are wellpreserved, with a typical clast showing various stages ofdecomposition from the core of the unweathered clast to theoutermost weathered rind. The rinds vary in thickness from a
few millimeters up to 1 cm and consist of fine-grained
actinolite, biotite, albite, and magnetite.
Near the biotite granite and dikes in the northern part ofthe outcrop area, the Mundy Breccia has been texturally
and mineralogically altered by alkali metasomatism. Alter-7
ation of basalts involved conversion of plagioclase laths tomicrocline and replacement of groundmass by muscovite,
biotite, and microcline. No evidence of intersertal or amygdaloidal textures is seen in these rocks. The skarn zone at
the contact of the Mundy Breccia with the Castner Marble
consists mostly of massive garnet (andradite) and scapolitewith minor amounts of calcite, biotite, magnetite, diopside(?), and actinolite. Isolated euhedral garnets attain sizes
up to 5 cm.
Equilibrium mineral assemblages indicate the Mundy
Breccia experienced at least low-grade metamorphism.
Total pressure at the time of metamorphism was at least
0.3 kb based on Harbour's (1972) estimates of thestratigraphic thickness of Precambrian rocks overlying theMundy Breccia.
Conclusions
The source of basalt that now makes up the clasts in the
Mundy Breccia could be from two basalt dikes that lie
unconformably below the breccia. However, it is not possible
to tell if the breccia formed from erosion of a lava flow orfrom shallow emplacement of a sill. Intersertal andamygdaloidal textures are common in lava flows as well as
high-level synvolcanic sills and dikes (Best, 1982). The lack ofsorting and stratification in the Mundy Breccia suggests thatthe breccia moved as a thick, viscous mudflow capable ofcarrying clay- to boulder-size clasts. There is no evidence toindicate if the Mundy Breccia is a single- or multiple-flow
unit, or whether the breccia traveled as subaerial or
subaqueous flow. Deposition of the breccia onto carbonate
sediments resulted in disruption of layering, folding, anddevelopment of intraformational breccias. The breccia wassubsequently exposed to weathering for some length of time,
as shown by spheroidally weathered clasts at various stages ofdecomposition in the upper portion of the volcanic mud-flow
breccia. In addition to weathered, sand-sized and smaller
basalt particles, the matrix was probably a mud composed ofminerals such as clay and Fe-oxides derived from weatheringof the basalt. Locally, quartz, perhaps a result of latersilicification, was abundant enough to cement the frameworkand matrix. Deposition of sediments of the Lanoria Formation,and sediments and volcanic rocks of the Thunderbird Group on top of the Mundy Breccia, was followed
by a period of plutonic igneous activity with intrusion ofseveral phases of the Red Bluff Granite complex. Emplacement of the high-level granite plutons resulted in stoppingof roof rocks and contact metamorphism of all pre-existingrocks.

References
Best, M. G., 1982, Fundamentals of igneous and metamorphic petrology:
W. H. Freeman & Co., New York, 656 pp.
Denison, R. E., and Hetherington, E. A., Jr., 1969, Basement rocks in far
west Texas and south-central New Mexico: New Mexico Bureau of
Mines & Mineral Resources, Circular 104: 1-16.
Dye, J. L., 1970, Geology of the Precambrian rocks in the southernFranklin Mountains, El Paso County, Texas: Unpublished MS thesis,
University of Texas at El Paso, 81 pp.
Harbour, R. L., 1960, Precambrian rocks at North Franklin Mountain,
Texas: American Association of Petroleum Geologists, Bulletin, 44:
1785-1792.
Harbour, R. L., 1972, Geology of the northern Franklin Mountains, Texas
and New Mexico: U.S. Geological Survey, Bulletin 1298, 129 pp.
Hoffer, R. L., 1976, Contact metamorphism of the Precambrian Castner
Marble, Franklin Mountains, El Paso County, Texas: Unpublished MS
thesis, University of Texas at El Paso, 77 pp.
McAnulty, W. N., Jr., 1967, Geology of the Fusselman Canyon area,
Franklin Mountains, El Paso County, Texas: Unpublished MS thesis,
University of Texas at Austin, 79 pp.
McCutcheon, T. J., 1982, Petrology and geochemistry of the Precambrian Red
Bluff Granite complex, northern Franklin Mountains, El Paso County,
Texas: Unpublished MS thesis, University of Texas at El Paso, 177 pp.
Ray, D. R., 1982, Geology of the Precambrian Red Bluff Granite complex,
Fusselman Canyon area, Franklin Mountains, El Paso County, Texas:
Unpublished MS thesis, University of Texas at El Paso, 297 pp.
Richardson, G. B., 1909, Description of the El Paso quadrangle, Texas:
U.S. Geological Survey, Geological Atlas, Folio 166, 11 pp.
Thomann, W. F., 1980, Petrology and geochemistry of the PrecambrianThunderbird Formation, Franklin Mountains, El Paso County, Texas:
Unpublished PhD dissertation, University of Texas at El Paso, 162 pp.
Thomann, W. F., 1981, Ignimbrites, trachytes, and sedimentary rocks of
the Precambrian Thunderbird Group, Franklin Mountains, El Paso,
Texas: Geological Society of America, Bulletin, 92: 92-100.
Thomann, W. F., and Hoffer, J. M., 1985, Petrology and geochemistry of
the Precambrian Mundy Breccia, Franklin Mountains, Texas: Texas
Journal of Science, 36: 267-281.
Thomann, W. F., and Hoffer, J. M., 1989, Stratigraphy, sedimentary structures, and petrology of the Middle Proterozoic Castner Marble, Franklin
Mountains, far west Texas: University of Wyoming, Contributions toGeology 27: 33-39.

Summary of the geochemistry and petrology of Proterozoic granitic rocks, Franklin Mountains, west Texas
W. M. Shannon and C. G. Barnes
Introduction
Middle Proterozoic granitic and associated volcanic rocks
are exposed for a distance of 25 km along the Franklin
Mountains, a westward-tilted Basin and Range fault block(Richardson, 1909; Harbour, 1972; Lovejoy, 1975) that extends northward from the center of El Paso, Texas (Fig. 1).
The granitic rocks intruded a Proterozoic metamorphic sequence (calc-silicates, basaltic breccia, and clastic rocks;
Harbour, 1960) and presumably cogenetic felsic volcanic
rocks (Thunderbird Group; Thomann, 1981).
The granitic rocks yield radiometric ages in the range 1.031.15 Ga (Wasserburg et al., 1962; Muehlberger et al., 1966;
Denison and Hetherington, 1969; Norman et al., 1987;

Copeland and Bowring, 1988). Coeval felsic activity occurred in Colorado (Pikes Peak; Hedge, 1970) and central
Texas (Enchanted Rock; Walker, 1988), and mafic activity
occurred in west Texas (Nellie layered mafic intrusions;
Hoover et al., 1985; Keller et al., 1989).
The granitic rocks have been inferred to be anorogenic(A-type) granite with a probable crustal source (Anderson,1983) or to be the product of fractionation of tholeiitic
magma typical of continental convergent margin and associated back-arc tectonic environments (Condie, 1982;
Norman et al., 1987). One of the objectives of our researchis to clarify the tectonic setting of these rocks by gaining amore complete understanding of their petrogenetic history.

Field setting and petrography
The sequence of intrusion of the granitic rocks is listed
below, from oldest to youngest (also see stratigraphic
section in Anthony et al., this volume).
Stage 1 is composed of porphyritic biotite—amphibolealkali-feldspar—quartz syenite (Fusselman Canyon sill) andporphyritic alkali-feldspar granite (sill in Edgar Park Quarry).
The sill exposed in Fusselman Canyon intruded the lower
member of the Lanoria Quartzite, and is up to 660 m thick.
The central part is alkali-feldspar—quartz syenite with hypidiomorphic granular texture and abundant glomerocrysts ofperthitic alkali-feldspar from 0.25 to 3.0 cm across. Ferrohedenbergite (Fe-rich clinopyroxene) occurs as anhedral
inclusions in the alkali-feldspar and as cores in amphibole.
Quartz phenocrysts reach as much as 2 mm in diameter.
Amphibole, biotite, quartz, alkali-feldspar, ilmenite, apatite,and zircon are interstitial. The margins of the sill are 1-5 m
thick and consist of porphyritic granite with euhedral quartz
and alkali-feldspar phenocrysts as much as 10 mm across.
Within 1 m of the contact, amphibole and biotite increase in
grain size and abundance. Fluorite occurs as

inclusions in the amphibole and interstitially (with clustersof acicular biotite). Zircon is a microphenocryst phase.
The porphyritic alkali-feldspar granite sill that crops out
in the vicinity of Edgar Park Quarry is also intruded into
the Lanoria Quartzite. The sill has porphyritic hypidiomorphic granular texture with euhedral quartz and alkalifeldspar phenocrysts. Biotite and amphibole clusters occur
in the groundmass. The biotite and amphibole areintergrown with fluorite and subhedral oxides. Zircon
microphenocrysts are also present.
Stage 2, the main intrusive phase, is biotite—amphibolealkali-feldspar granite. It forms about 80% of the Franklin
Mountain granitic rocks (Fig. 1). It is hypidiomorphic granular and contains perthitic alkali-feldspar with cores of subhedral plagioclase (as calcic as An 36 ) or albite. Plagioclase,quartz, and zircon are interstitial, as are intergrowths ofamphibole with biotite, albite, quartz, zircon, and fluorite.
Ilmenite occurs as lamellar zones in biotite and as discrete
subhedral grains in amphibole. Prismatic apatite occurs as
inclusions in amphibole and biotite. Magnetite is sparse andits abundance increases as the ratio of amphibole to biotite
decreases.
Stage 3 consists of biotite—amphibole—alkali-feldspar—
quartz syenite. It is abundant in the southern end of theFranklin Mountains (Fig. 1). Stage-3 rocks are similar in
texture and mineralogy to the Stage-2 granite and contain
deuteric grunerite(?, after fayalite?). Quartz-syenite dikes
contain alkali-feldspar phenocrysts with plagioclase cores(An 23 to An 31 ). In these dikes, poikilitic amphibole and biotite
enclose alkali-feldspar and acicular apatite. Zircon forms
abundant euhedral prisms as well as acicular crystals in the
groundmass.
Stage-4 leucocratic biotite—alkali-feldspar granite underlies Granite Peak on the northeastern edge of Fusselman
Canyon (Fig. 1). It is also well exposed in Tom Mays Park,
where it intrudes volcanic rocks of the Thunderbird Group.
It consists of an hypidiomorphic granular arrangement ofquartz and alkali-feldspar with interstitial biotite, fluorite,
ilmenite, and magnetite. Alkali-feldspar phenocrysts arepresent locally.
Stage 5 consists of biotite—arfvedsonite—alkali-feldspar
granite. This rock type has two textural varieties: one isporphyritic and the other is equigranular. In the porphyritic
variety, subhedral quartz and alkali-feldspar (now microcline) are phenocryst phases and arfvedsonite, astrophyllite,
fluorite, and zircon form glomerocrysts. The matrix is afine-grained hypidiomorphic granular arrangement ofquartz, albite, and microcline. Some quartz and microcline
crystals are broken and have poikilitic overgrowths withinclusions of matrix quartz and alkali-feldspars. Groundmass zircon is poikilitic; opaque oxides are rare.
Some dikes and sills of the arfvedsonite granite are banded(i.e. line rock). The bands consist of flow-foliated laths ofarfvedsonite. Streaks within the layers have poikilitic arfvedsonite, biotite, quartz, and zircon. The upper contact of somesills is pegmatitic, with well-developed comb layering.
An abandoned tin-mining site is located north of Mundy
Springs. East—west-trending pegmatitic ore veins cut Stage-2
granite. The margins of these veins are leucocratic ferrimuscovite—topaz—alkali-feldspar granite. The rock is finegrained, hypidiomorphic granular, and forms sharp intrusive
contacts with the host granite. Perthitic alkali-feldspar, quartz,
ferrimuscovite, and topaz occur with accessory fluorite, zircon, ilmenite, and magnetite.

Mineral compositions
Alkali-feldspar compositions range from Ab 58 An 4 Or 38 in
the quartz syenites to Ab54An 1 Or 45 in the most voluminous
alkali-feldspar granites. Plagioclase cores in alkali-feldspar
phenocrysts are Ab66An 32 Or 2 in the quartz syenites andAb 62 An 36 Or 2 in the main-phase granite.
Biotite Fe/(Fe + Mg) ranges from 0.9 to 1.0 (annite); theratio increases from quartz syenite to main-phase granite.
Mn content remains approximately constant and Ti content
decreases with increasing Fe/(Fe + Mg). Biotite in thearfvedsonite granite is Al-deficient and contains excess Si.
In general, Cl fills about 10% of the hydroxyl site but fluorine was rarely detected. However, the ratio F/(F + Cl +H 2 0) ranges from 0.5 to 0.75 in annite in the arfvedsonite
granite (and in ferrimuscovite in the topaz granite).
Amphibole compositions (Fig. 2) range from ferroedenite
to ferroactinolite when the Fe+3/Fe+2 ratio is calculated
according to Speer and Kimble (1981). This ratio ranges
from 0.03 to 0.17. Edenite substitution appears to be thedominant mechanism of compositional variation, but Ti
Tschermak's substitution is significant in amphiboles fromthe least siliceous quartz syenites. Ti content decreases and
Mn content increases as Fe/(Fe + Mg) increases. The ratioof halogens to hydroxyl in ferroedenite is about 0.25 andthe fluorine content of the ferroedenitic amphiboles decreases with increasing Fe/(Fe + Mg) and Mn/(Mn + Ti). In
addition to the magmatic amphiboles, interstitial deutericgrunerite is locally present.

Major-element, trace-element, andisotopic compositions
The chemical compositions of these rocks are typical of
metaluminous alkali granites (e.g. Fig. 3, Table 1). Majorand trace-element trends are continuous from quartz syenite
to granite and plot along linear trends. Si, Cs, Rb, Be, Hf, Y,
Th, and the rare-earth elements (REE) increase in thissequence, whereas Zr, Ti, P, Fe, Mg, Mg/(Fe + Mg), Mn,
Sc, Al, K, Na, Ba, and Sr decrease. Table 1 shows that therocks have high contents of high-field-strength elements (Zr,Hf, Nb, and Ta), total alkalis, and REE.
REE-element patterns are approximately parallel and total
REE abundance increases from the quartz syenite to the
granitic compositions (Fig. 4). The Eu anomaly increases
with differentiation. Although the main-phase granite is texturally variable, its REE concentrations are all similar. The

alkali-feldspar crystals) which indicate partial crystallization
at pressures higher than that of intrusion (Nekvasil, 1990).
Several lines of evidence suggest that the oxidation statef(0 2 ) in the syenitic and granitic magmas was low. These
include the extreme Fe/(Fe + Mg) of mafic minerals andwhole rocks, the mineral assemblage ferrohedenbergite +ferroedenite + annite + ilmenite, and the presence ofarfvedsonite as a magmatic phase (Wones, 1981, 1989).
The hedenbergite + ilmenite assemblage may be represented by the reaction: 3 sphene + 2 magnetite + 3 quartz =3 ferrohedenbergite + 3 ilmenite + 0 2 . Rocks on thehedenbergite + ilmenite side of the buffer require an
oxygen fugacity at or below the fayalite—quartz—
magnetite buffer (Wones, 1989).
The presence of fluorite as a late magmatic phase in mostrock types, as well as the moderate fluorine content of latestage ferromagnesian minerals, indicates that both fluorine
and water were dissolved in the magmas. However, the lackof abundant pegmatitic rocks suggests that fluid saturation
was probably reached at near-solidus conditions.
Discussion
REE patterns of the Fusselman Canyon quartz syenite are
identical to those of the early granitic sill near Edgar Park
Quarry (Table 1), an indication that the two sills may represent a single intrusive event.
The quartz syenites have whole-rock δ18O values of about8‰ and quartz values of about 9‰. δ18O of the main granite
is about 7‰ (whole rock) and 8‰ (quartz). These values
suggest that hydrothermal alteration did not lower thewhole-rock δ18O value. The initial 87Sr/86Sr for theFranklin Mountain granites is about 0.7034 (Norman et al.,1987). This value is slightly higher than that of the uniform
reservoir at 1.14 Ga (0.7032). ENd for the granites and latealkali-basalt dikes ranges from +3.8 to +1.4 at 1.14 Ga(Patchett and Ruiz, 1989; Norman et al., 1987).
Conditions of emplacement
A pseudoternary plot of normative albite, quartz, andorthoclase indicates that the rocks differentiated from the
Ab—Or join toward the 0.5 kbar water-saturated minimum(Fig. 5). This is consistent with pressure of emplacement
estimated on the basis of stratigraphic reconstruction of about
1 kbar. Low-pressure emplacement was preceded by partial
crystallization in a deeper magma chamber because the compositional fields for the bulk of the quartz syenites andgranites lie in a zone that is compatible with crystallization atpressures closer to 3 kbar. This inference is consistent withthe subsolvus nature of the rocks (plagioclase cores in

Mineralogical and isotopic data suggest that magmas parental to the Franklin Mountain granitic suite were derived
from a source that was relatively reduced, had mantle-like Srand Nd-isotopic compositions, but oxygen-isotope compositions that were slightly elevated relative to the normal
mantle. Trace-element abundances indicate that the parental
magma was derived from a source that was enriched in highfield-strength elements or that a more mafic HFSE-richmagma underwent protracted fractionation to produce thetrace-element enrichments observed in the granitic suite. We
prefer the latter explanation (i.e. an alkali-basaltic parent)
because of the difficulty of attaining such high HFSE
concentrations in a metaluminous magma by partial melting.
This difficulty is caused by the fact that in a crustal-source
region, the HFSE are held in refractory accessory phases thathave low solubility in metaluminous magmas. Sparse, latestage, Si0 2 -undersaturated mafic alkalic dikes cut the granitic
rocks and indicate that such compositions were present.
Major- and trace-element variations suggest that differentiation among the granitic rocks could have occurred byprogressive removal of alkali-feldspar, ferroedenite, annite,
ilmenite, apatite, and zircon (Fig. 6). In addition, the Rb/ Y
ratio of the suite is nearly constant (Fig. 7), as would be
expected if fractional crystallization was the dominant process. Fractional crystallization cannot explain the higher
18O value of the syenitic rocks relative to the granites. Thishigher value indicates that the granitic rocks cannot be related to the syenites by the simple process of fractional
crystallization. In spite of this isotopic discrepancy, wetested the possibility that a mafic syenitic magma wasparental to the granitic compositions.
Major-element mass-balance calculations (Bryan et al.,1969) were used to model fractional crystallization ofmafic quartz-syenitic magma to make successively moredifferentiated granitic daughter magmas. Alkali-feldspar,
ferrohedenbergite, ferroedenite, annite, ilmenite, andapatite were the minerals removed in the models. The
results of the calculations are combined with trace-element
data in Fig. 6. Trace-element variations (as a function of
calculated fraction of liquid remaining; F) indicate thatfractional crystallization can produce the range of granitic
compositions observed.

12
LeMone, D. V. (eds.), Border stratigraphy symposium: New Mexico
Bureau of Mines & Mineral Resources, Circular 104: 1-16.
Dye, J. C., 1970, Geology of the Precambrian rocks in the SouthernFranklin Mountains, El Paso County, Texas: Unpublished MS thesis,
University of Texas at El Paso, 81 pp.
Giret, A., Bonin B., and Leger, J. M., 1980, Amphibole compositional
trends in oversaturated and undersaturated alkaline plutonic ring complexes: Canadian Mineralogist, 18: 481-495.
Harbour, R. L., 1960, Precambrian rocks at North Franklin Mountain,
Texas: American Association of Petroleum Geologists, Bulletin, 44:
1785-1792.
Harbour, R. L., 1972, Geology of the northern Franklin Mountains, Texas
and New Mexico: U.S. Geological Survey, Bulletin 1298: 129 pp.
Hedge, C. E., 1970, Whole rock Rb-Sr age of the Pikes Peak batholith,
Colorado: U.S. Geological Survey, Professional Paper 700-B: 86-89.
Hoover, J. D., Keller, G. R., and Hills, J. M., 1985, The Nellie intrusion: abasic stratiform intrusion in the Central basin platform in west Texas:
EOS, 66: 1136.
Keller, G. R., Hills, J. M., Baker, M. R., and Wallin, E. T., 1989, The
Nellie intrusion: a basic stratiform intrusion in the Central basin
platform in west Texas: Geology, 17: 1049-1052.
Lovejoy, E. M. P., 1975, An interpretation of the structural geology of theFranklin Mountains, Texas; in Seager, W. R., Clemons, R. E., andCallender, J. F. (eds.), Las Cruces country: New Mexico GeologicalSociety, Guidebook 26: 261-268.
McAnulty, W. N., Jr., 1967, Geology of the Fusselman Canyon area,
Franklin Mountains, El Paso County, Texas: Unpublished MS thesis,
University of Texas at Austin, 79 pp.
McCutcheon, T. J., 1982, Petrology and geochemistry of the Precambrian
Red Bluff Granite complex, Fusselman Canyon area, Franklin Mountains, El Paso County, Texas: Unpublished MS thesis, University ofTexas at El Paso, 177 pp.
Muehlberger, W. R., Hedge, C. E., Denison, R. E., and Marvin, R. F.,
1966, Geochronology of the midcontinent region, United States: Part 3,
Southern area: Journal of Geophysical Research, 71: 5409-5426.
Nekvasil, H., 1990, Reaction relations in the granite system:
Implications for trachytic and syenitic magmas: American
Mineralogist, 75: 560571.

Nelson, L. A., 1940, Paleozoic stratigraphy of Franklin Mountains, west
Texas; in De Ford, R. K., and Lloyd, E. R. (eds.), West Texas-New
Mexico Symposium: American Association of Petroleum Geologists,
Bulletin, 24: 157-172.
Norman, D. I., Condie, K. C., Smithy, R. W., and Thomann, W. F., 1987,
Geochemical and Sr and Nd isotopic constraints on the origin of thelate Proterozoic volcanics and associated tin-bearing granites from theFranklin Mountains, west Texas: Canadian Journal of Earth Sciences,24: 830-839.
Patchett, P. J., and Ruiz, J., 1989, Nd isotopes and the origin of Grenvilleage rocks in Texas: Implications for Proterozoic evolution of theUnited States mid-continent region: Journal of Geology, 97: 685-695.
Ray, R. D., 1982, Geology of the Precambrian Red Bluff Granite complex,
Fusselman Canyon area, Franklin Mountains, El Paso County, Texas:
Unpublished MS thesis, University of Texas at El Paso, 295 pp.
Richardson, G. B., 1909, Descriptions of the El Paso quadrangle, Texas:
U.S. Geological Survey, Atlas, Folio 166, 11 pp.
Streckeisen, A., 1974, Classification and nomenclature of plutonic rocks:
Recommendations of the IUGS Subcommission on the Systematics of
Igneous Rocks: Geologischen Rundschau, 63: 773-786.
Taylor, H. P., Jr., and Sheppard, S. M. F., 1986, Igneous rocks: I. Processes of isotopic fractionation and isotope systematics; in Valley, J.
W., Taylor, H. P., Jr., and O'Neil, J. R. (eds.), Stable isotopes in hightemperature geological processes: Mineralogical Society of America,
Reviews in Mineralogy, 16: 227-271.
Thomann, W. F., 1981, Ignimbrites, trachytes and sedimentary rocks of
the Precambrian Thunderbird Group, Franklin Mountains, El Paso,
Texas: Geological Society of America, Bulletin, 92: 94-100.
Walker, N. W., 1988, U-Pb zircon evidence for 1305-1231 Ma crust in the
Llano uplift, central Texas (abs.): Geological Society of America,
Abstracts with Programs, 20: A205.
Wasserburg, G. J., Wetherill, G. W., Silver, L. T., and Flawn, P. T., 1962,
A study of the ages of the Precambrian of Texas: Journal ofGeophysical Research, 67: 4021-4047.
Wones, D. R., 1981, Mafic silicates as indicators of intensive variables in
granite magmas: Mining Geology, 31: 191-212.
Wones, D. R., 1989, Significance of the assemblage titanite + magnetite +quartz in granitic rocks: American Mineralogist, 74: 744-749.

Day 2: The Potrillo volcanic field
J. M. Hoffer, W. Chen, V. T. McLemore, and E. Y. Anthony
Driving distance and time: 135.5 mi, 8 hrs, 4 stops.
Road-log starting point: Intersection of Mesa and Doniphan, just west of I-10 (Mesa exit)
in northwest El Paso.
Mileage

0.0 Stoplight at Mesa and Doniphan, straight ahead on
Texas Farm Road 260 (Country Club road). 0.6 0.6
Stoplight; continue straight. 0.4
1.0 Stoplight; continue straight. 0.5
1.5 Cross Rio Grande. 0.5
2.0 Stop sign; enter New Mexico on NM-184, straightahead. 0.6
2.6 Stop sign; turn right onto NM-273. 1.5
4.1 Prepared food plant on right. 0.2
4.3 Turn left onto Airport Road. 0.2
4.5 Santa Teresa School on right. 2.3
6 . 8 Tu r n l e f t . 0 . 4
7.2 Stop sign; cross railroad tracks and bear right on
A-17, continue parallel to railroad tracks on the
La Mesa surface. The La Mesa surface represents
the upper surface of the Rio Grande rift basin-filldeposits; here it is a dune sand composed of subrounded, clear to iron-coated quartz grains (70%)

and subordinate opaques, zircon, tourmaline, andpyroxene (Hoffer, 1976). 0.8
8.0 Turn left on A-03. 0.8
8 . 8 C a t t l e gu a rd . 5 . 9
14.7 Turn left and then right onto the old railroad bedof the El Paso and Southwestern Railroad,continue west. 3.0
17. 7 C at tl e gu a rd. 1.4
19.1 East Potrillo Mountains at 2:00-3:00. These mountains(Fig. 1) exhibit Laramide, early Tertiary, and late
Tertiary deformation, and contain Permian, Lower
Cretaceous, and Cenozoic strata (Seager and Mack,
1986). 2.6
21. 7 C at tl e gu a rd. 5.1
26. 8 C at tl e gu a rd. 0.4
27.1 West Potrillo Mountains at 12:00-2:00; Mt. Riley and
Mt. Cox at 2:00; and East Potrillo Mountains at
2:00-3:00. The West Potrillo Mountains (Fig. 1) area broad topographic high covered by late Tertiary
and Quaternary alkaline lavas. The Potrillo volcanic
field as a whole is of an alkaline nature and varies,
with differentiation, from basanites through alkali
basalts to trachybasalts (Table 1, Fig. 2). The
plateau rises 120 to 240 m above the La

Mesa surface and is dotted by numerous volcanoes
that reach elevations of nearly 1675 m (Hoffer andSheffield, 1981). Mantle and crustal xenoliths andxenocrysts are common in the flows and maars of
the area and include spinel peridotite, amphibole,and plagioclase (Ortiz, 1979). Abundant crystals offeldspar occur as loose crystals or as inclusions inbombs associated with the older cones. The crystals
are anhedral to euhedral, white to colorless, andrange from several millimeters to nearly 5 cm indiameter. They range in composition from anorthoclase to plagioclase (Hoffer and Hoffer, 1973).
Cinder cones are the most abundant volcanoes in
the range; over 150 cones have been mapped (Hof-fer, 1976). They range from 60 to 150 m in heightand 300 to 900 m in diameter. They appear to be ofat least two ages, with the younger cones large andsteep-sided, with slopes from 20 to 25°, and older
cones more subdued, with slopes from 10 to 20°.
Two maars, Mt. Riley (not to be confused with thetrachydacite Mt. Riley, see mileage 72.5) and
Malpais, occur in the West Potrillo Mountains. They
have been described by Bersch (1977) and Page
(1975), respectively. 0.4
27.6 Caliche outcrops mark location of northwest-trending
fault. 4.9
32.5 Road to right leads to Mt. Riley ranch; continue
straight. 0.5
33.0 Caliche outcrops mark location of north-striking
fault. 0.5
33.5 Basalt lava flows of the West Potrillo Mountains at
1:00-3:00. 0.9
34.4 A-4 to the left; continue straight. 3.9
38.3 Camel Mountain, a local landmark at 12:00. 0.7 39.0
Cattle guard; Guzman Lookout, a young cinder cone at
1:30. 0.7
39.7 Turn right onto dirt road and drive north to Malpais
maar. Caution, drive slowly as the road is rough.
1.4
41.1 Stop I: Leave cars and walk about 100 yards down
the arroyo to examine the air-fall and base-surge
deposits on the east side of Malpais maar (Fig. 3).
Eruption began at Malpais maar with deposition of
up to 15 m of bedded debris over the surroundingarea. Thickest accumulations of tuffs are locatedwhere the base-surge clouds encountered nearby
cinder cones and were forced to move upslope. The
deposits wrap gently around this topography, dippingaway from the cone summits at 5 to 20°. The style oferuption later changed to lava fountaining and tephra
production, with construction of a cinder cone andflow complex within the crater. Late-stage activity
included extrusion of parasitic lava flows andintrusion of dikes into the bedded-tuff sequence. The
dikes are concentrated along the maar rim. A small,maar-type vent is found along the strike continuationof one of the dikes (Page, 1975). Sedi-

15
mentary structures, typical of base-surge deposits(Crowe and Fisher, 1973), are present throughout
the bedded-tuff sequence. These structures include:(1) long wavelength, low amplitude, undulatory
bedding (plane-parallel lamination); (2) asymmetrical, ripple-like to dune-like bed forms; (3) small-tolarge-scale crosslaminations; and (4) bedding-sag
deformation under bombs and blocks. The maar rim
is nearly circular, approximately 1250 m in diameter, and is perhaps 10 to 15 m in height, although
the lower contact is not exposed. Inward-dipping
beds are exposed at three locations along the maarrim, dipping 30 to 35° towards the center of thecrater. Outer-flank beds near the rim dip 5 to 20°
away from the vent. Farther from the rim, beddingin the tuffs generally conforms to subdued pre-tuff
topography, although crosslaminations in ripple anddune bedforms indicate a consistent, radial transport
for the base-surge clouds. 1.0
42.1 Stop II: Cinder quarry in the crater of Malpais
maar. Return to cars and retrace route to A-17. 2.4
44.5 Turn left on A-17 and travel east. 6.9
51.4 Mt. Riley ranch road on left; continue straight.
5.8
57.2 Cattle guard. 1.5
58.7 Turn left on A-08. 3.3
62.0 Site of abandoned marble quarry at the base of theEast Potrillo Mountains at 9:00. 2.0
64.0 Cattle guard. 2.8
66.8 Turn right onto A-14. 1.6
68.4 Turn left and travel north to Hunt's Hole. 0.2 68.6
Turn left and travel around the south end of Hunt'sHole. 0.1
68.7 View of Hunt's Hole. Hunt's Hole is the smallest of
three circular maar volcanoes in the westernmost
part of the La Mesa surface. The other maars are

Kilbourne Hole located 3.4 km north of Hunt'sHole, and Potrillo maar 16.5 km to the south. The
three maars are aligned north-south along the inferred down-to-the-west Fitzgerald fault; this fault
is not exposed in the maars or on the La Mesa
surface but is expressed by a line of topographic
depressions. Hunt's Hole is excavated in mudstoneand sandstone of the Pleistocene Santa Fe Groupand an overlying basalt flow which represents thesouth end of the Aden-Afton volcanic centers. The
age of the maar-forming event is probably lessthan 50,000 yrs based on the small degree ofdissection and lack of weathering of the base-surge
tuffs (Stuart, 1981). 1.0
69.7 Cattle guard. 0.9
70.6 Small spatter cone at 12:00. 1.8
72.4 Stop III: Kilbourne Hole overlook (lunch stop).
Kilbourne Hole (Fig. 4) is a maar 3.3 km long and 2
km wide. The floor of the depression is 250 to 90 m
below the surface of the surrounding desert and isbordered by a 20 to 170 ft rim. Recent work by
Seager (1987) indicates that an important process inthe formation of the present crater is caldera-likecollapse at the end of the volcanic cycle. Evidence forthis interpretation includes: (1) the presence of large
downfaulted masses of the tuff-ring material locatedat the base of the innter crater wall and beneath thecrater floor, and (2) the small volume of pyroclastic
material in the maar rim and on the La Mesa plains ascompared to the volume of the crater. Age estimatesof the maar range from 180,000 yrs (Seager et al.,1984) based on a K-Ar date on post-maar lavas at
Potrillo maar, to 24,000 yrs based on the amount of
pedogenic-carbonate development in soils of theKilbourne rim ejecta (Gile, 1987). In addition, afission-track date on apatite obtained

16from the outer surface of a basalt bomb yields anage of 80,000 yrs (V. Harder, pers. comm. 1988).
0.1
72.5 Return to main road; bear right toward Mt. Riley. Mt.
Riley and Mt. Cox are two peaks from asingle pluton. Mt. Riley, to the northeast, is thelarger and reaches an elevation of 5915 ft (1804m). The pluton is a steep-sided, domal intrusion ofdense trachydacite (nomenclature of LeBas et al.,1986) of probable Tertiary age. It intrudes Cretaceous(?) sedimentary rocks and Tertiary volcanic andsedimentary rocks, and is composed of 58%
plagioclase (andesine), 20% potassium feldspar, 14%
quartz, 7% biotite and hornblende, and 1% magnetite(Millican, 1971). It contains 65.2% SiO 2 and 7.9%
K 2 O + Na 2 O. 1.6
74.1 Turn right. 2.0
76.1 A-12 to the right; continue straight. 1.0
77.1 Abandoned ranch on right. 0.2
77.3 Gardner cones (Afton flows) at 12:00. 0.6
77.9 Cinder cone of the West Potrillo field (two tanks
on top of the cone) at 12:00. 1.1
79.8 Cross flows of the West Potrillo field. 0.4
79.4 Gate; continue straight. 0.4
79.8 Turn left at gate. 0.1
79.9 Turn right at gate. 0.9
80.8 Road is on the flows of the West Potrillo field, Adenflows on the right at 12:00-5:00. 1.5
82.3 Gate. 3.1
85.4 Beyond water tank, Aden Crater at 2:00. 0.3
85.7 Gate; bear right and pass through a second gate andthen left. 1.8
87.5 Turn right toward Aden Crater. 0.6
88.1 Stop IV: Aden Crater, which lies on top of the Adenflows and along the Aden rift. It is thought that the
Aden flows were emplaced along fissures associatedwith the Robledo fault and that the crater was built
from younger central eruptions. It is classified as asmall shield volcano or lava cone (DeHon, 1965).
The base of the cone is approximately 4.5 km indiameter and composed of thin pahoehoe lava flows
which slope gently outward from the summit at 35°(Hoffer, 1976). The cone rises about 50 m above thesurrounding terrane and covers an area of approximately 15 km', with an estimated minimum lava
volume of 0.3 km3 (Kahn, 1987). The age of thecrater is probably less than 100,000 yrs based on
surface freshness and soil development (tile,
1987).
The formation of Aden Crater involved the following sequence of events. A series of thin basaltic
flows issued from a central vent and spread out
radially, building a small shield volcano. This period of quiet activity was followed by a more explosive gas phase in which spatter was ejected,
constructing a spatter rampart around the outer rimof the crater. The magma, after degassing, issued
once again from the vent and formed a lava lake
which was contained by the spatter rim. At several
locations, lava breached the rim and flowed down
the flank of the volcano. The low viscosity of thelava is indicated by the fact that several overflowing
lava tongues and tubes are only 12-24 cm wide, butflowed 40-50 m down the side of the volcano.

At the close of this effusive period of activity, aportion of the ponded lava drained back down thevent and the lava lake solidified and collapsed overthe vent. After solidification of the lava lake, which
produced tension cracks parallel to the spatter rim, aperiod of minor fountaining occurred in the central
crater and on the southeast flank of the cone producing a series of spatter cones. Final activity wasconfined to a fumarole on the east rim of the crater(Hoffer, 1990).
Geologic tour of Aden Crater
Distance: roundtrip = 2.5 km, time = 3 hrs, 5 stops.
Total
Distance0

Parking area, northwest flank of Aden Crater (Fig.
5). 360 m
0.36 km Walk up the slope of the volcano toward the southsoutheast to view a series of spatter cones at Stop
1. Spatter cones are generally steep-sided, small
volcanic cones which are composed of ejected clots
or globs of magma that were still somewhat fluidwhen they accumulated. Note the taffy-likeappearance on the outer surface of the fragments.
Note also the alignment of the cones toward thesouthwest along a probable fracture or fissure zone
on the flank of Aden. 300 m
0.66 km Climb up over the spatter rim and across the floor of
Aden Crater to the large spatter cone at Stop
2. The floor of Aden was formed by a series ofbasaltic lava flows that flooded the interior crater
and were dammed by the spatter rim. The depthof this former lava lake is unknown, but a hole
drilled into the lake in 1970 by NASA penetrated
over 30 m of lava. The spatter cone of this stopsits on top of the lava lake and was one of thelatest events in the development of Aden. On thewest side of the cone (Fig. 5) note that a portionof the crater rim has slid some distance down the17slope. The grooved and flattened surface upslope
from the slide block indicates that the rim had
not yet completely solidified when the slide
occurred. 180 m
0.84 km Walk south and then southwest around the west sideof the crater vent to Stop 3. This is the location of
the primary vent of the volcano; it measuresapproximately 160 by 100 m and is over 17 m
deep. The present configuration of the vent is due
to collapse which occurred at the close of theperiod of lava-lake formation. During its active
phase, the vent alternated between fountaining andquiet effusion of basaltic lava. Note, across thevent to the southwest, the moderate-size hill covered with red to reddish-brown spatter. This spatterprobably accumulated downwind during fountaining episodes of the vent. 270 m
1.10 km Continue south from the vent and then follow thecrude trail on the spatter rim of the crater to thefumarole at Stop 4. The Aden fumarole is situatedin the southeast rim of the crater. The fumarole
opening is approximately 2 m in diameter at thesurface and is 33 m deep. In 1928, the well preserved remains of a ground sloth (Nothrotheriops)
were discovered at the bottom of the fumarole.
C'4 data on the sloth bones indicate an age of
11,080 ± 200 yrs. 300 m
1.40 km Walk east down the flank of the volcano to viewthe collapsed portion of a lava tube at Stop 5. The
collapsed portion of the lava tube measures 230 mlong, 60 m wide, and 9 m deep. A small spattermound has been constructed at the southeast rimof the tube. This lava tube can be traced under therim of Aden, just north of the fumarole, into thefloor of the crater where numerous collapse
depressions and tube cross-sections exist.
End of Aden Crater walking tour. Return bywalking the lower slopes of the volcano, approximately 1.1 km. 0.6
88.7 Return to main road and turn right. 4.2
92.9 Turn right on A-17 and travel parallel to railroad
tracks. 0.5
93.4 Cattle guard. 4.5
97.9 Cattle guard. 2.5
100.4 Cross basal flows of the Afton flows. 0.5
100.9 Turn left, cross railroad tracks and bear left. 0.2
101.1 Cattle guard. 0.1
101.2 Turn right onto B-4. 2.1
103.3 Pass ranch. 0.1
103.4 Bear left and continue straight. 1.3
104.7 Government installation on right. 0.7
105.4 Cattle guard. 2.4
107.8 Turn right onto B-8. 1.0
108.8 Cattle guard and start of pavement. 0.4
109.2 Cattle guard, El Paso Natural Gas pumping stationon left. 1.7
110.9 Cattle guard. 1.1
112.0 Little Black Mountain cinder cone at 9:00. 1.7
113.7 Turn off to Black Mountain on right.
Four eruptive centers, extending north-south approximately 38 km2, occur along the eastern marginof the La Mesa surface. They are, from north tosouth, Santo Tomas, San Miguel, Little Black
Mountain, and Black Mountain. Black Mountainand Santo Tomas both display multiple eruptions.
The Black Mountain area consists of eight volcanic cones and at least six flows (Hoffer, 1971).
In the center of the area are four cinder cones, thelargest being the Black Mountain cone. It is symmetrical in outline, with a diameter of 700 m and aheight of 30 m.
Cinder-spatter cones are located north and southof Black Mountain. These cones possess a lowsloping base of cinder and spatter and are cappedby a rim of agglutinated spatter. The layers of
spatter and driblet dip steeply inward toward thecenter of the cone. The cones range from 30 to 80
m in diameter and from 5 to 14 m in height. 2.5
116.2
Cattleguard.
1.9
118.1
Cattleguard.
0.4
118.5
Cross
bridge. 0.8 119.3 Intersection of
NM-28 and B-8; turn left onto NM28, which is part of the Don Juan de Oñate trail, ahistoric route commemorating Don Juan de Oñate's
travels through New Mexico. Cotton, alfalfa, hay,and chili are grown in the fields along NM-28. 0.3
119.6 Sierra Vista nursery and greenhouses on the left.
Organ Mountains at 1:00-2:00. On Day 3 we will
drive along the foothills of this mountain range.
1.8
121.4 Road junction with County Road 189 at stop sign.
Turn left to remain on NM-28. 1.1
122.5 Bend in road. Little Black Mountain at 10:00. 0.3
122.8 Edge of La Mesa, one of numerous small farming
communities on NM-28. La Mesa is Spanish for"table" and was named for the nearby Black Mesa
lava flow. 2.6
125.4 La Clinica de Familia on the right. Colorful murals
such as this are common on public buildings in Old
Mexico. Edge of San Miguel. 0.6
126.0 Junction with County Road 192. Continue straighton NM-28 through San Miguel. 0.2
126.2 San Miguel Catholic Church on the left is built withboulders from the San Miguel and Santo Tomasflows to the west. 0.3
126.5 North end of San Miguel. San Miguel lava flow is at
9:00. This flow has been dated by K/Ar as 490,000± 30,000 yrs (Seager et al., 1984). However, recent
sedimentologic studies by Gile (1990) suggest thatthe age is approximately 200,000 yrs. The Santo
Tomas flow, consisting of a cinder cone and threeassociated flows (Hoffer, 1971), is at 10:00- 11:30. 0.5
127.0 MP 20. Entering Stahmann Farms, one of the world's
largest pecan groves which markets the Del Cerro
pecans. Stahmann Farms end at the Rio Grande,although many smaller, privately owned groves
continue along NM-28. 3.0
130.0 MP 23. Crossing the Rio Grande. Picacho Peak andthe Robledo Mountains at 11:00. On Day 3 we will
discuss the geology of the Robledo Mountains(Hawley, 1978). 3.1
133.1 Entering San Pablo. 1.6
134.7 Junction with County Road 373. Continue straighton NM-28. 0.8
135.5 Road junction. Turn left onto Calle de Parian into La
Mesilla (Spanish for "Little Table"). La Mesilla
was built on a small mesa above the Rio Grande in181850 and has been a state monument since 1957.
The Butterfield Overland Trail passed through La
Mesilla and famous historic characters such as Billy
the Kid visited the town. Today a plaza surroundedby the church and numerous shops and restaurants
draw tourists from throughout the world.
End of road log.
Geochemistry of late Cenozoic lavasof southern New Mexico
Introduction
Late Cenozoic mafic lavas from southern New Mexico
fall into two chemical groups: (a) low-silica, nephelinenormative basanites, alkali basalts, and trachybasalts, and(b) higher-silica, hypersthene-normative subalkali basalts(Fig. 2). Both types have similar ages and both are found in
the axis of the rift as well as on its flanks. In any individual

volcanic field the lavas will belong predominantly to one
chemical group or to the other. For instance, the Jornadadel Muerto, Animas, and Carrizozo fields belong to thehypersthene-normative group, and the Geronimo, ElephantButte, and Palomas fields belong to the nephelinenormative group (Anthony et al., in review). The Potrillo
volcanic field is of the nepheline-normative type, withlavas from various parts of the field ranging from basanites
through alkali basalts to trachybasalts (Table 1, Fig. 2).
Inter-field chemical variations
The hypersthene-normative group of lavas is characterized
by lower concentrations of the incompatible elements Rb, Sr,
Zr, Nb, and TiO 2 (Figs. 6a, 7a). Although the factors
responsible for generating the two chemical groups, e.g.
volatile contents, degree of partial melting, and depth ofmagma generation, are still under investigation, the data at
hand suggest that variable degrees of partial melting may
have played a significant role. Higher concentrations ofincompatible elements (particularly Nb and Sr) in the nepheline-normative lavas are consistent with their genesis bylesser degrees of partial melting. Also, as noted by Luhr et19al. (1989) for similar lavas from San Luis Potosi, Mexico,of the various hypotheses for generating undersaturated lavas, incremental partial melting best explains the observed
correlation
betweensilicaundersaturation
andincompatible-element enrichment.
Intra-field chemical variations
Magmas of both the nepheline and hypersthene groups
appear to have resided temporarily in low-pressure reservoirs. Lavas of both groups share a similar phenocryst population of olivine and plagioclase with minor clinopyroxene.
These phases are near liquidus at pressures less than approximately 10 kb, corresponding to depths at or above the
Moho (30 to 35 km) in this region (Keller et al., 1990).
Further, within a volcanic field, the incompatible elements
Zr, Nb, Sr, and Rb increase up-section, suggesting sufficient
residence time for crystal fractionation and/or crustal assimilation. This increase in incompatible elements up-section
has been documented for lavas from the Jornada del Muerto
field (a hypersthene-normative field) and the Potrillo volcanic field (Table 2, Figs. 6b, 7b). The shallow level of
ponding is consistent with barometric determinations forcrustal and mantle xenoliths transported with the magmas
during their final ascent.

Acknowledgments
Chemical compositions reported herein were obtained by
XRF analysis at the New Mexico Bureau of Mines & Mineral Resources. We thank Chris McGee, manager of the

XRF facility, for his expert and patient analysis. This manuscript benefited from the comments of Nancy McMillan,
New Mexico State University, and Jacques Renault, New
Mexico Bureau of Mines & Mineral Resources.
References
Anthony, E. Y., Hoffer, J. M., and Waggoner, W. K. (in review), Compositional diversity in late Cenozoic lavas in the Rio Grande rift andBasin and Range province, southern New Mexico.
Bersch, M. G., 1977, Petrography and geology of the southern West
Potrillo Basalt Field, Dona Ana County, New Mexico: Unpublished
MS thesis, University of Texas at El Paso, 60 pp.
Crowe, B. M., and Fisher, R. V., 1973, Sedimentary structures in basesurge deposits with special reference to cross-bedding, Ubehebe
Craters, Death Valley, California: Geological Society of America,
Bulletin, 94: 663-682.
DeHon, R. A., 1965, Maare of La Mesa: New Mexico GeologicalSociety, Guidebook 16: 204-209.
Gile, L. H., 1990, Chronology of lava and associated soils near San
Miguel, New Mexico: Quaternary Research, 33: 37-50.
Gile, L. H., 1987, A pedogenic chronology for Kilbourne Hole, southern
New Mexico. Soil Science Society of America, Journal 51: 746-760.
Hawley, J. W. (compiler), 1978, Guidebook to Rio Grande rift in New
Mexico and Colorado: New Mexico Bureau of Mines & Mineral Resources, Circular 163: 241 pp.
Hoffer, J. M., 1990, Geologic tour of Aden Crater: El Paso GeologicalSociety, Guidebook to the Fall 1990 Field Trip, 42 pp.
Hoffer, J. M., 1976, Geology of Potrillo basalt field, south-central New
Mexico: New Mexico Bureau of Mines & Mineral Resources, Circular
149: 30 pp.
Hoffer, J. M., 1971, Mineralogy and petrology of the Santo Tomas—
Black Mountain basalt field, Potrillo volcanics, south-central New
Mexico: Geological Society of America, Bulletin, 82: 603-611.
Hoffer, J. M., and Hoffer, R. L., 1973, Composition and structural state offeldspar inclusions from alkali olivine basalt, Potrillo basalt, southern
New Mexico: Geological Society of America, Bulletin, 84: 2139-2142.
Hoffer, J. M., and Sheffield, T. M., 1981, Geology of the West Potrillo
Mountains, south-central New Mexico: El Paso Geological Society,
Guidebook, pp. 79-82.
Kahn, P. A., 1987, Geology of Aden Crater, Dona Ana County, New
Mexico: Unpublished MS thesis, University of Texas at El Paso, 90 pp.
Keller, G. R., Morgan, P., and Seager, W. R., 1990, Crustal structure,gravity anomalies, and heat flow in the southern Rio Grande rift andtheir relationship to extensional tectonics: Tectonophysics, 174: 21-37.
LeBas, M. J., Le Maitre, R. W., Streckeisen, A., and Zanettin, B., 1986, A
chemical classification of volcanic rocks based on the total alkali-silica
diagram: Journal of Petrology, 27: 745-750.
Luhr, J. F., Aranda-Gomez, J. J., and Pier, J. G., 1989, Spinel—lherzolitebearing Quaternary volcanic center in San Luis Potosi, Mexico. I. Geology, mineralogy and petrology: Journal of Geophysical Research, 94:
7916-7940.
Millican, R. S., 1971, Geology and petrography of the Tertiary Riley—
Cox Pluton, Doña Ana County, New Mexico: Unpublished MS thesis,
University of Texas at El Paso, 87 pp.
Ortiz, T. S., 1979, Megacrysts and mafic and ultramafic inclusions of thesouthern West Potrillo Basalt Field, Dona Ana County, New Mexico:
Unpublished MS thesis, University of Texas at El Paso, 95 pp.
Page, R. 0., 1975, Malpais Maar volcano: New Mexico Geological Society, Guidebook 26: 135-137.
Seager, W. R., 1987, Caldera-like collapse at Kilbourne Hole maar, New
Mexico: New Mexico Geology, 9: 69-73.
Seager, W. R. (in press), Geology of southwest quarter of Las Cruces andnorthwest El Paso 1° x 2° sheets, New Mexico: New Mexico Bureau of
Mines & Mineral Resources, Geologic Map 60, scale 1:125,000.
Seager, W. R., and Mack, G. H., 1986, Summary of geology of East
Potrillo Mountains area: El Paso Geological Society, Guidebook, pp.
158-159.
Seager, W. R., Hawley, J. W., Kottlowski, F. E., and Kelley, S. A., 1987,
Geology of east half of Las Cruces and northeast El Paso 1° X2º 2°
sheets, New Mexico: New Mexico Bureau of Mines & Mineral
Resources, Geologic Map 57, scale 1:125,000.
Seager, W. R., Shafiqullah, M., Hawley, J. W., and Marvin, R. F., 1984,
New K—Ar dates from basalts and the evolution of the southern Rio
Grande rift: Geological Society of America, Bulletin, 95: 87-99.
Stuart, C. J., 1981, Hunt's Hole Maar volcano, Doña Ana County,
south-central New Mexico: El Paso Geological Society, Guidebook,pp. 6472.

20
Day 3: Las Cruces to AlbuquerqueG. R. Keller and V. T. McLemore
The third day of the trip will follow the itinerary set
forth in the following guidebook: Callender, J. F., Hawley,
J. W., and Keller, G. R., 1989, Rio Grande rift: 28th International Geological Congress, Guidebook to Field Trip
T318, 37 pp. A copy of this guidebook will be provided aspart of the registration package. In addition to discussions
by the field-trip leaders, Richard W. Harrison and Richard
P. Lozinsky will help with the stops in the Truth or Consequences area, and Lee R. Russell and Richard P.
Lozinsky will help with stops in the Albuquerque basin.

21

New Mexico Bureau of Mines & Mineral Resources Bulletin 137, April 1991

Open-system magmatic evolution of the Summer Coon and Del Norte
volcanoes, Conejos Formation, San Juan Mountains volcanic field,
Colorado
Don F. Parker, D. Anne Grau, and Charles D. Thomas
Department of Geology, Baylor University, Waco, Texas 76798-7354

Introduction
Many modern petrologic investigations have emphasized
open-system behavior in magmatic systems, where crystal
fractionation, assimilation of country rock, and magma mixing provide genetic relations between end members of igneous suites (Hildreth et al., 1986; Dungan et al., 1986;
Smith and Leeman, 1987). Petrologists have, of course,
recognized these controls for many years, but usually were
only able to document fractional crystallization because oflimitations of technique. Today, electron-probe microanalysis of minerals, trace-element analysis, and isotopic studiesof rocks aid petrologists in identifying, respectively, out-ofrange phenocrysts in mixed rocks and the geochemical andisotopic signatures of assimilated country rock in magmas.
This study tests the applicability of open-systembehavior to the evolution of igneous rocks of the Summer
Coon and Del Norte volcanoes. Field evidence for opensystem behavior, including compositional banding of lava,
mafic enclaves, and inclusions of basement gneiss in dikes
and lavas are examined, and major- and trace-element data
bearing on the question of assimilation of crustal rocks arepresented. The first description of the Del Norte volcano isgiven, and data from previous studies of the Summer Coon
volcano are amplified.
Geologic setting
The San Juan volcanic field (SJVF) was formed during
the mid-Cenozoic as part of a broad continental arc associated with shallow subduction of oceanic lithospherealong the southwestern margin of North America (Lipman
et al., 1971; Cross and Pilger, 1978). The SJVF is anerosional remnant of a larger field that covered much ofsouth-central Colorado and northern New Mexico (Steven,1975). Subduction-related magmatism spanning theinterval from 35 to 30 Ma produced a coalescing
assemblage of composite volcanoes of largely andesitic
composition, which is overlain by thick ash-flow tuffs
erupted from caldera sources during the period 30-26 Ma(Lipman et al., 1970; Lipman, 1989). Volcanism associatedwith development of the Rio Grande rift began 26 Ma ago
and has continued to the present. The San Luis Valley,
formed by an asymmetric graben associated with the Rio
Grande rift, lies east of the SJVF.
Summer Coon and Del Norte volcanoes
The Summer Coon and Del Norte volcanoes belong tothe older episode of intermediate volcanism that is represented in the eastern SJVF by the Conejos Formation (Fig.
1). The Summer Coon volcano (SCV) is a highly dissectedcomposite cone 12 km in diameter, with a well-exposed,
extensive radial dike swarm (Lipman, 1968, 1976; Mertzman, 1971; Noblett and Loeffler, 1987). Igneous rocks associated with the SCV span the compositional spectrum
from high-potassium basaltic andesite to rhyolite (Fig. 2).
The Del Norte volcano (DNV) is a highly eroded silicic
shield composed of thick lava flows, which span a relatively
narrow range of compositions from high-potassium silicic
andesite to dacite.

The oldest eruptive unit of the SCV is the Tsa unit of
Lipman (1976) (Ta in Fig. 1). It is largely composed ofthick, coarse volcanic breccias that are mostly debris-avalanche and lahar deposits formed in association with explosive eruptions. Lava flows are locally common within the
Ta unit, but are subordinate in volume to the breccias. Rockcompositions range from basaltic andesite to dacite, with
andesite most abundant. The Ta unit is overlain by packagesof silicic lavas (Ts in Fig. 1), mostly on the southern flankof the volcano, erupted in the sequence rhyolite—dacite(Thomas, 1988; Parker and Thomas, 1988). Eruptions at
SCV were fed from a large radial dike swarm. Dike compositions mirror the compositions of flows. Basalticandesite and andesite dikes are mostly thin (0.5-3 m thick),whereas dacite and rhyolite dikes can be much thicker. A
central region of fine- to coarse-grained intrusions probably
originally underlay the summit region of the composite
cone (Mertzman, 1971).
Whole - rock chemistry and mineralogy
Chemical analyses of igneous rocks from the SCV havepreviously been given in Mertzman (1971) and Zielenski andLipman (1976). Conclusions of this report are based on 60
new whole-rock major- and trace-element analyses completed at Baylor University. Broadly, the igneous rocks constitute a suite ranging from high-potassium basaltic andesite
to rhyolite, with an average composition of silicic andesite(Figs. 2, 3). These results are similar to those reported by
Mertzman (1971) for the SCV and by Lipman et al. (1978)
for early intermediate-composition rocks of the entire SJVF.
Harker diagrams (not illustrated) show linear correlations
between oxides of major elements and silica, with only thealkali oxides showing appreciable enrichment.
Major-phenocryst chemistry of SCV rocks is summarized
in Figs. 3 and 4. Plagioclase (pc), magnetite (mt), and apatite
occur in all samples, and zircon occurs in most samples.
Olivine (ol) occurs only in basaltic andesite and someandesite. Clinopyroxene (cpx) is abundant in basaltic
andesite and occurs with orthopyroxene (opx) in andesite.

Some dacites have opx. The only alkali-feldspar grains discovered are suspected to be disaggregated crystals from
inclusions of hypabyssal intrusions in a highly porphyritic
silicic-andesite lava. DNV rocks are similar, but containabundant hornblende (hbl) in all rocks, including andesite,where it occurs in-lieu of pyroxene. Hornblende is only
abundant in dacite and some rhyolite of the SCV. Biotite (bi)is common in the most evolved rocks of both volcanoes.
Summary of trip itineraryThis field trip consists of seven stops. The first four stops
are concerned with the Summer Coon volcano (SCV). The
stops provide a representative look at the different geologic
units constituting the SCV (Fig. 5). Stop 1 examines intermediate lava flows of the Ta unit. Stop 2 emphasizes dikes,
especially an andesite dike with abundant, large gneiss inclusions. Stop 3 focuses upon the breccias within the Ta unit,
and Stop 4 examines late silicic flows of the SCV. Afterlunch, the last three stops concern the Del Norte volcano(DNV). Stop 5 examines the large dacitic dike of Indian Head.
Stop 6 is an overview of a silicic eruptive center on the west
flank of the DNV, and Stop 7 examines intermediate lava anda dacite dome on Lookout Mountain, just outside of Del
Norte, Colorado. The general theme of all the stops will be tounderline field evidence for open-system behavior in these
volcanoes.

Field trip
Directions to Stop 1
The trip starts at the intersection of US-285 and US-160
in downtown Alamosa, Colorado. Proceed west on US-160
through Monte Vista, Colorado, to Del Norte, Colorado(about 30 mi). In Del Norte, turn north on CO-112 andproceed approximately 3.2 mi to the intersection with the La
Garita Road (Rio Grande County 33) (Fig. 5). Turn left onto
the La Garita Road, which becomes unpaved. The roadenters a strike valley between west-facing cuestas formed
by, respectively, the Fish Canyon Tuff to the west and theCarpenter Ridge Tuff to the east (both lumped in Ty of Fig.
1). These tuffs were erupted from calderas in the central San
Juan Mountains. Continue north on the La Garita Road tothe intersection with Forest Service Road 660 (about 6 mifrom the intersection with CO-112). Turn left on Forest
Road 660 (also known as Road A32) and proceed about 1
mi to Stop 1, which is located along the eastern extremity of
the Summer Coon volcano where a small stream valley
emerges from the eroded remnants of the volcano.
Stop 1The Ta unit of Fig. 1 is dominated by coarse, crudely
layered, poorly sorted breccia deposits, which largely formedby mass-wasting processes on the flanks of the SCV (Fig. 6).
Lava flows, however, are locally abundant in the Ta unit,

and Stop 1 contains an unusual concentration of these lava
flows in the upper part of the Ta unit.
At least three dark, vesicular andesite lavas are overlain
by two flows of silicic andesite. The andesite lavas (Analysis
A, Table 1) crop out in the small hill next to the road, and in
streamcuts, and possess highly scoriaceous flow tops. Silicicandesite lava (Analysis B, Table 1) forms cliffs along thenorthern margin of the streambed. The two silicic-andesite
flows are similar chemically and mineralogically and could
be lobes from a single eruption. The upper silicic flow is 1520 m thick, overlies andesitic breccia, and is capped by a thinflow breccia. Well-developed foliation is parallel to the 12°
dip of the unit near the base, but steepens upward into ramp
structures.
Andesite lavas, such as exposed at this stop, show
textural and mineralogical evidence for magma mixing andassimilation of crustal rocks. Evidence for magma mixing
includes out-of-range phenocrysts, compositional bandingof lavas, and mafic enclaves with acicular hornblende interpreted to be quickly chilled lumps of mafic magma mixed
with andesite magma (Bacon, 1986). Out-of-range phenocrysts include olivine mantled by cpx and highly resorbed
opx mantled by cpx in andesite (Fig. 7). Normally, neitherof nor opx would be present in a mafic andesite. Evidence

for assimilation includes small inclusions in lava of granite
and granite gneiss, some of which have been partly assimilated
juding from their amoeboid shapes, and large xenoliths of
biotite gneiss in an andesite dike. These features will be
examined and discussed at various stops on our trip.
Stop 2
From Stop 1, proceed west on FS Road 660
approximately 3.1 mi to the junction with FS Road 659,
which is a spur leading to the La Ventana area. Travel
north on FS-659 about 1.7 mi and park near La Ventana, anatural arch formed by weathering of a large dacitic dike.
Walk west-northwest about 200 m from La Ventana to thebiotitegneiss-xenolith locality.
Stop 2A consists of an andesite dike 2 m in diameter(Analysis C, Table 1), which contains angular to rounded
inclusions of biotite gneiss (Analysis D, Table 1) up to 1 m
in diameter (Fig. 8). The andesite contains pc, cpx, opx, andmt microphenocrysts and, other than the xenoliths, isunremarkable. The inclusions are composed of alkali feldspar, plagioclase, biotite and minor opaque minerals. No
quartz was observed, although one sample was quartz-normative. The other two analyzed samples had minor normative nepheline. Most of the alkali feldspar is now
sanidine due to heating from the andesite, although
appreciable microcline remains. Several of the inclusions
contain small rounded blebs of brown glass (in somedevitrified), indicating incipient melting of the inclusions.
The andesitic magma certainly was hot enough to melt thegneiss. A temperature of 1030 ± 50°C was estimated fromcoexisting pyroxene compositions from a similar Summer
Coon andesite (Parker and Thomas, 1988).
Previous workers have documented the importance ofassimilation of locally heterogeneous crust to the evolutionof the Conejos magma series (Colucci et al., 1989; Lipman
et al., 1978). Crustal influences can be modeled through theuse of incompatible-element ratios (Davidson et al.,1987). In particular, the K/Rb ratio for Conejos rocks is agood measure, because both are very incompatible in themagmas, and no mineral, except plagioclase to a smalldegree and hornblende, seems capable of fractionating K
relative to Rb. Hornblende, although abundant in silicic
andesite of the DNV, is absent from andesite and basaltic
andesite of the SCV. The K/Rb ratio should, therefore, remain constant in a sequence dominated by fractional crystallization, or, possibly, might decrease somewhat.
Fig. 9 shows K/Rb ratios of SCV and DNV rocks plottedagainst Rb as an index of differentiation. The overall trend
could be interpreted as a result of simple fractional crystallization, or coupled assimilation—fractional crystallization, with typical high Rb, low K/Rb crustal assimilant.
Many andesites, however, and some dacites from SCV and
DNV have K/Rb values higher than those of basaltic andesite. In fact, most andesite analyses fall along mixing
lines between basaltic andesite and a field represented by
the three analyzed gneiss inclusions. This suggests
evolution of the SCV and DNV magmatic suites by a twofold process: (1) assimilation of crustal material represented
by the gneiss inclusions (about 25 wt% is indicated bycomparison of Rb and Ba values), then (2) fractional
crystallization to yield the more evolved dacites andrhyolites. This model, of course, does not preclude an
earlier episode of deep crustal assimilation by basalt tocreate basaltic andesite. We hope to further test the
proposed model by additional geochemical work.

lanche and lahar deposits resulting from mass wasting off
the steep flanks of the composite cone of the SCV following
explosive eruptions, are well exposed at Stop 3. Many beds
show inverse grading of clasts (Fig. 10). Clasts within thebeds include strongly compositionally banded rock, breadcrust bombs, and radially jointed blocks (Fig. 11). The latter
indicate that many of the clasts were still hot when incorporated in the lahars. The clasts range in composition frombasaltic andesite (Analysis G, Table 1) to low-silica dacite.
Proceed west on Forest Road 660 less than 0.5 mi to thejunction with the Old Women Creek Road. Turn left (south)and go approximately 2.8 mi to Stop 4. Pull to the side of
the road next to a faded brown signpost.
Stop 2B will allow examination of the large dacite dikethat forms La Ventana. This dike is about 6 m thick andcontains pc, hbl, bi, opx, cpx, and mt phenocrysts. The
margins of the dike have horizontal columnar joints. Piecesof a vitrophyre from this dike can be found along the eastern
flank of the dike. Unfortunately, we have not located thesource of the vitrophyre. A two-pyroxene model estimate forthe temperature of the dacite dike was 970 ± 50°C.
Return back to FS Road 660, then turn right (west).
Proceed west on FS Road 660 approximately 5 mi to thecrest of a low divide to reach Stop 3.
Stop 3
Stop 3 will allow the examination of typical exposures of
the Ta unit of Fig. 1, and a small andesitic dike (Analysis F,
Table 1). Coarse breccias, probably largely debris-ava-

26
Stop 4Late in its eruptive history, the SCV produced silicic lava
flows (Ts in Fig. 1), particularly along its southern margin.
At this stop, we will examine rhyolitic and dacitic lavas
belonging to the Tsr and Tsd units of Lipman (1976). In thissection, partly brecciated and silicified, cream-colored,sparsely porphyritic rhyolite is overlain by platy, darkbrown, porphyritic dacite. The rhyolite ranges up to 150 m
in thickness and contains only pc, bi, and mt phenocrysts. A
chemical analysis (Analysis H, Table 1) from this locality
shows 74.7 wt% SiO2. Lipman (1976) reported a range
from 71 to 73%, and Zielinski and Lipman (1976) publishedan analysis from this locality with 71.1% SiO 2 . The
overlying dacite lava, unfortunately, is not suited foranalysis, but a similar sample taken northeast and updip
from Stop 4 had 63.7% SiO 2 .
One of the major differences between the SCV and the
DNV is the presence of a minor rhyolite component in the
SCV but not in the DNV. Zielinski and Lipman (1976)
presented trace-element data that supported a fractional crystallization origin of rhyolite from andesite for the SCV.
Parker and Thomas (1988) came to similar conclusions, andalso proposed that these final eruptions from SCV were theproducts of a compositionally zoned magma chamber below
SCV, in which rhyolite overlay denser, highly porphyritic
dacitic magma. We are currently investigating the origin ofrhyolite at SCV in more detail.
Continue southeast on FS Road 660 to the junction with
the North River Road. Turn left (east) and go about 3.7 mito the junction of the North River Road with CO-112. Go
south (right) 0.2 mi to the Del Norte Centennial Park forlunch. After lunch, return to the North River Road andproceed approximately 3 mi from the intersection with CO112 to Stop 5, the Indian Head dike.
Stop 5This report is the first formal description of the Del
Norte volcano (Grau, 1989). Lipman (1976) mentioned a
volcano near Del Norte as a source for some of the Tcv
units on his map. Previously, Mertzman (1971) mapped the
Del Norte rocks as part of the Summer Coon volcano.
The Del Norte volcano (DNV) is a small silicic shield
volcano with a diameter of about 6 km. At one time itprobably resembled the larger, much younger, and better
preserved San Antonio volcano of the Taos Plateau. The
DNV is composed largely of thick hornblende-andesite andhornblende-dacite lava flows (many of them 100-200 m
thick). Coarse breccias underlie flows; these breccias are inpart autobreccias and in part debris-avalanche and lahar
deposits. Lavas were erupted from a central dike swarm,
much of which is covered today by alluvium from the Rio
Grande (Fig. 1).
Stops 5, 6, and 7 examine vent areas of the DNV. Stop 5
examines the Indian Head dike (Analysis I, Table 1), a
north—south dike which fed a shallow intrusion andprobable dacite dome along the north margin of the DNV(Fig. 12). Several silicic dikes in the SCV and DNV havebulbous extensions or terminations. These dike flarings
have been interpreted by Moats and Fink (1987) as theresult of decreased lithostatic load on the dikes along themargins of the volcanoes.
No contact between the dike and country rock is visible
near the highway. However, the western margin of the dikeis brecciated and some of the clasts are vitrophyric, sug-gesting proximity to the wall. The dike contains a numberof inclusions of hornblende andesite, hypabyssal igneousrocks, and hornblende-rich metamorphic rocks.
Return to the town of Del Norte via the North River Road
and CO-112. Turn right (west) onto US-160 at the intersection with CO-112. Go west about 3.7 mi and stop on thenorthern shoulder of the highway about 100 m past aroadcut in saddle.
Stop 6
Stop 6 is an overview of a dacite eruptive center along
the western margin of the DNV (Fig. 13). The majority of

with the hornblende andesite (Fig. 18). We interpret thesestreaks as smeared-out mafic enclaves. Also present aresubrounded inclusions up to 6 cm in diameter ofplagioclase—hornblende rock, some with acicular
hornblende, as well as rounded inclusions of granite or
granite gneiss. The inclusions with acicular hornblende orother evidence for quick cooling may also represent mafic
enclaves. If the process of mingling of magmas illustrated by
the lava at this stop had proceeded further, most of theevidence for mixing would have disappeared. We suggestthat some of the compositional banding present in andesitic
rocks of the SCV could have resulted from such mixing.
the shield of the DNV (Tds in Fig. 13) is composed ofhornblende andesite and dacite (SiO 2 range 58-65 wt%).
The Tds unit forms a northwest-dipping cuesta north of the
highway in Fig. 13. Three west-northwest-striking silicic
dikes cut the Tds unit in exposures along the north bank of
the Rio Grande. These three dikes appear to be radial to acommon center south of the words "Rio Grande" in Fig.
13. All of these probably fed flows, but only the central
dike can be traced upwards into a silicic-dacite lava (68.3
wt% SiO2)• The southwesternmost dike (about 10 m thick)can easily be seen from US-160 (Figs. 14, 15). These lavas
and dikes are the most evolved rocks in the DNV(Analyses J and K, Table 1). The discovery of these dikes
and other eruptive centers such as Lookout Mountainestablishes the DNV as an independent eruptive center
separate from the slightly younger SCV.
Return to Del Norte via US-160. At the western marginof Del Norte, turn south onto the Pinos Creek Road.
Proceed 0.25 mi to Stop 7A. Park in the abandoned sandquarry on the southeast side of road.
Stop 7A
The Tds unit of the DNV is characterized by thick (usuallymore than 100 m) flows of silicic andesite and dacite (Fig.
16) that appear to have flowed several kilometers. This stop
examines a hornblende-andesite lava (Analysis L, Table 1) of
the Tds unit that shows evidence for mixing with a more
mafic magma (Fig. 17). The flow at this stop overlies a flow
breccia several meters thick, which in turn rests upon a
breccia of probable laharic origin.
This flow contains streaks over 1 m long and several
centimeters thick of vesicular, aphanitic mafic rock mingled

References

Stop 7B
Return to the vehicles and drive up the gravel road topark next to a Del Norte water reservoir. Walk about 250
m southeast along a footpath.
Lookout Mountain (with the large white "D" on it) is theerosional remnant of a dacite dome of the DNV (Analysis
M, Table 1) (Fig. 19) similar chemically to the silicic-dacite
flows and dikes of Stops 5 and 6. The base of the dome ismarked by a prominent vitrophyre several meters thick,which dips to the southeast at 30°. The crest of Lookout
Mountain contains ramp structures indicating flow from thesoutheast (Fig. 19). The probable vent for the dome is along
the southeastern margin of Lookout Mountain. The dome
appears to have cut upwards through the Tds unit at a shallow angle (before tilting associated with the development of
the San Luis graben steepened the dip). Its original extent
was most likely not very different from the present outcrop.

Acknowledgments
This work was supported by the Department of Geology,
Baylor University, and the University Research Council of
Baylor University. The senior author benefited from a Summer Sabbatical from Baylor University awarded to pursue
this research. We would also like to thank the landowners of
the area for permission to work on their properties. We areindebted to many former Baylor University students for their
help in collecting and processing samples, and especially to
Richard Furner, Gary Henderson, and Greg Elkins. The
excellent mapping by Peter Lipman served as a base formuch of this project.

Bacon, C. R., 1986, Magmatic inclusions in silicic and intermediate volcanic rocks: Journal of Geophysical Research, v. 91, pp. 6091-6112.
Colucci, M. T., Dungan, M. A., Lipman, P. W., and Moorbath, S., 1989,
Petrology and isotope geochemistry of the Conejos Formation, SE San
Juan volcanic field: implications for multiple parent magmas andcrustal interactions (abs.): New Mexico Bureau of Mines & Mineral
Resources, Bulletin 131, p. 57.
Cross, T. A., and Pilger, R. H., Jr., 1978, Constraints on absolute motion
and plate interaction inferred from Cenozoic igneous activity in thewestern United States: American Journal of Science, v. 278, pp. 865902.
Davidson, J. P., Dungan, M. A., Ferguson, K. N., and Colucci, M. T.,
1987, Crust—magma interaction and the evolution of arc magmas: theSan Pedro—Pellado volcanic complex, southern Chilean Andes:
Geology, v. 15, pp. 443-446.
Dungan, M. A., Lindstrom, M. M., McMillan, N. J., Moorbath, S., Hoefs,
J., and Haskin, L. A., 1986, Open system magmatic evolutionof the Taos Plateau Volcanic Field, northern New Mexico, I. The petrology and geochemistry of the Servilleta Basalt: Journal ofGeophysical Research, v. 91, pp. 5999-6028.
Grau, D. A., 1989, Comparative petrologic evolutions of the Del Norteand Summer Coon volcanoes, eastern San Juan Mountains, Colorado:
Unpublished BS thesis, Baylor University, Waco, Texas, 115 pp.
Hildreth, W., Grove, T. L., and Dungan, M. A., 1986, Introduction to
special section on open magmatic systems: Journal of Geophysical Research, v. 91, pp. 5887-5889.
Lipman, P. W., 1968, Geology of the Summer Coon volcanic center,eastern San Juan Mountains, Colorado: Colorado School of Mines
Quarterly, v. 63, pp. 211-236.
Lipman, P. W., 1976, Geologic map of the Del Norte area, eastern San
Juan Mountains, Colorado: U.S. Geological Survey, Miscellaneous Investigations Series Map 1-952.
Lipman, P. W., 1989, Oligocene—Miocene San Juan volcanic field, Colorado; in Chapin, C. E., and Zidek, J. (eds.), Field excursions to volcanic
terranes in the western United States, Volume I: Southern Rocky
Mountain region: New Mexico Bureau of Mines & Mineral Resources,
Memoir 46, pp. 303-305.
Lipman, P. W., Prostka, H. J., and Christiansen, R. L., 1971, Cenozoic
volcanism and plate-tectonic evolution of the western United States, I.
Early and middle Cenozoic: Royal Society of London, PhilosophicalTransactions, v. 271, pp. 217-248.
Lipman, P. W., Steven, T. A., and Mehnert, H. H., 1970, Volcanic history of
the San Juan Mountains, Colorado, as indicated by potassium—argon
dating: Geological Society of America, Bulletin, v. 81, pp. 2329-2352.
Lipman, P. W., Doe, B. R., Hedge, C. E., and Steven, T. A., 1978,
Petrologic evolution of the San Juan volcanic field, southwestern Colorado: Pb and Sr isotope evidence: Geological Society of America, Bulletin, v. 89, pp. 59-82.
Mertzman, S. A., Jr., 1971, The Summer Coon volcano, eastern San Juan
Mountains, Colorado; in James, H. L. (ed.), Guidebook of the San Luis
Basin, Colorado: New Mexico Geological Society, Guidebook 22, pp.
265-272.
Moats, W., and Fink, J., 1987, Geometry of dikes at Summer Coon
volcano, south-central Colorado (abs.): Geological Society of America,
Abstracts with Programs, v. 19, no. 7, p. 775.
Noblett, J. B., and Loeffler, B. M., 1987, The geology of Summer Coon
volcano near Del Norte, Colorado; in Beus, S. S. (ed.), Rocky
Mountain Section of the Geological Society of America, Field Guide,v. 2, pp. 349-352.
Parker, D. F., and Thomas, C. D., 1988, Petrologic evolution of theSummer Coon volcano, San Juan Mountains, Colorado (abs.): Geological Society of America, Abstracts with Programs, v. 20, no. 7, p. A71.
Peccerillo, A., and Taylor, S. R., 1976, Geochemistry of Eocene calcalkaline volcanic rocks from the Kastamonu area, northern Turkey: Contributions to Mineralogy and Petrology, v. 58, pp. 63-81.
Smith, D. R., and Leeman, W. P., 1987, Petrogenesis of Mount St. Helens
dacitic magmas: Journal of Geophysical Research, v. 92, pp. 1031310334.
Steven, T. A., 1975, Middle Tertiary volcanic field in the southern Rocky
Mountains: Geological Society of America, Memoir 144, pp. 75-94.
Thomas, C. D., 1988, Petrologic evolution of the Summer Coon
volcano, San Juan Mountains, southwestern Colorado: Unpublished BS
thesis, Baylor University, Waco, Texas, 89 pp.
Zielinski, R. A., and Lipman, P. W., 1976, Trace-element variations at
Summer Coon volcano, San Juan Mountains, Colorado, and the originof continental-interior andesite: Geological Society of America,
Bulletin, v. 87, pp. 1477-1485.

29

New Mexico Bureau of Mines & Mineral Resources Bulletin 137, April 1991

Geomorphic and neotectonic evolution along the margin of theColorado Plateau and Rio Grande rift, northern New Mexico
Mark A. Gonzalez' and David P. Dethier²'Department of Geology, University of New Mexico, Albuquerque, New Mexico 87131; 'Department of Geology, Williams College, Williamstown, Massachusetts 01267

Introduction
A number of investigators have examined fluvial systems
as a means of deciphering the geologic evolution of the Rio
Grande rift (e.g. Bryan, 1938; Bryan and McCann, 1936,
1938; Wright, 1946; Ruhe, 1967; Lambert, 1968; Bachman
and Mehnert, 1978; Hawley, 1978; Kelley, 1979; Manley,
1979a, b; Kelson, 1986a, b; Kelson and Wells, 1987; Wells etal., 1987; Dethier and Harrington, 1987; Dethier et al., 1988).
Ironically, the geomorphology of the lower Rio Chama, thelargest tributary of the upper Rio Grande, has not been
examined in detail. The lower reach of the Rio Chama seems
especially suited for geomorphic analysis for several reasons.
First, the valley of the Rio Chama and tributary arroyos
expose extensive remnants of geomorphic surfaces, providing
valuable spatial constraints on former base-level positions inthis portion of the Rio Grande rift. Second, someinvestigators (C. E. Stearns, pers. comm. 1987, 1990; S. G.
Wells, pers. comm. 1987) believe the Rio Chama may be theancestral headwaters of the Rio Grande system, antedating
the reach of the Rio Grande found upstream of the Rio
Chama/Rio Grande confluence. The upper reach of the Rio
Grande is believed to be geologically young, having formed
since the mid-Pleistocene (Wells et al., 1987). Third, thefortuitous juxtaposition of geomorphic features and radiometrically dated volcanic material provides temporalconstraints on the evolution of this portion of the Rio Grande
rift. Finally, the Rio Chama has experienced climatic, tectonic, and volcanic perturbations, all of which appear to haveleft a geomorphic signature on the landscape.
The purpose of this excursion along the margin of the Rio
Grande rift and Colorado Plateau of northern New Mexico is
two-fold: (1) to examine evidence of the long-term(approximately 8 my) evolution of the Rio Chama/Rio Grande
drainage system; and (2) to inspect neotectonic features thatchronicle the timing, magnitude, and kinematics of movements on rift-margin structures.
Geologic setting and previous investigations
The Rio Grande rift extends from central Colorado
through New Mexico and south to west Texas and
Chihuahua, Mexico, a distance of over 1000 km. The
southern portion of the rift is morphologically indistinct fromthe Basin and Range province, while the central and northern
portions, which are bounded to the west by the Colorado
Plateau and southern Rocky Mountains, are distinct morphotectonic basins (Baldridge et al., 1984). The central rift in
New Mexico consists of three asymmetric half-grabens (SanLuis, Espanola, and Albuquerque–Belen basins; Fig. 1).
Important structural features of the Espanola basin include
mountain-front bounding faults to the east along the Sangrede Cristo Mountains (Cabot, 1938; Menges, 1988; Vernon
and Riecker, 1989), the intra-basin NNE–SSW-trending
Embudo fault (Muehlberger, 1979; Aldrich, 1986; Aldrich

and Laughlin, 1984; also called the Santa Clara fault inHarrington and Aldrich, 1984), the intra-basin Velarde graben which parallels and is bounded in part by the Embudo
fault zone (Manley, 1976a, 1979a), the N–S-trending Pajarito fault zone (Los Alamos fault of Kelley, 1979) whichis the contemporary western boundary of the Espanolabasin (Gardner and House, 1987; Gardner and Goff, 1984;
Golombek, 1983), and N–S-trending ancient rift-margin
fault zones such as the Canada de Cochiti fault zone
(Gardner and Goff, 1984) and the Cannes fault zone
(mapped by Manley, 1982; Figs. 1, 2).
Recently completed seismic profiles and gravity measurements (Biehler et al., in press; W. S. Baldridge, pers.
comm. 1990; unpublished data from SAGE project—Summer of Applied Geophysical Experience) have supplemented
previous geophysical studies (e.g. Cordell, 1978, 1979, andreferences therein), enhancing our understanding of the subsurface geology of the Espanola basin. Geophysical evidence
suggests that the deepest parts of the Espanola basin are the
NE–SW-trending Velarde graben (Fig. 1) southeast of the
Embudo fault zone, where Tertiary rift-fill sediment andPaleozoic/Mesozoic sedimentary rocks are perhaps 23 km(about 7000-10,000 ft) thick, and the eastern side of the
Pajarito fault zone (Biehler et al., in press). In contrast, thenorth side of the Embudo fault zone has only 1 km thickness(less than 3500 ft) of rift-fill sediments (W. S. Baldridge,
pers. comm. 1990; unpublished seismic data from the 1990
SAGE course). The distinct change in sediment thickness
across the Embudo fault suggests that the Espanola basin is
structurally segmented, and we refer to the shallow
northwestern part of the basin as the Abiquiu embayment.
The rift-fill sediment and paleontology have been studied
by numerous investigators. Two seminal works, which provide comprehensive bibliographies, are the work by Galusha
and Blick (1971) and the recent publication by Ingersoll et al.
(1990). Fig. 3 provides a schematic view of the stratigraphy
along the Rio Chama and Rio Grande valleys from Ghost
Ranch to White Rock Canyon. Upper Paleozoic andMesozoic sedimentary strata are found on the Colorado Plateau, and Eocene through Quaternary rift-fill sediment andvolcanic rock dominate the strata of the Rio Grande rift.
Geomorphic features and late Cenozoic
drainage evolution
Evidence for two periods of aggradation, three periods ofvalley incision, and the formation of four pre-Quaternary andas many as 10 Quaternary surfaces occurs along the margin of
the Colorado Plateau, Abiquiu embayment, and eastern flankof the Jemez Mountains. We refer to the periods ofaggradation and incision as: (1) the period of late Miocene
through Pliocene incision; (2) the Puye aggradational period(Pliocene); (3) the pre-Bandelier incisional period (about 1.7
to 1.5 Ma), (4) the Bandelier aggradational period (1.5–1.1 Ma); and (5) the post-Bandelier incisional period (from1.1 Ma to the present). In the following section we describe
the evidence constraining the timing, spatial occurrence, andphysical characteristics of the periods of aggradation andincision, as well as of the pre-Quaternary and Quaternarysurfaces along approximately 90 km of the Rio Chama and
Rio Grande valleys.
For correlation of surfaces, we rely primarily on the heightof the surface above the axial-stream channel, as well as theradiometric ages of underlying or overlying (for buried
surfaces) deposits where available. Irregular patterns of erosion and burial of soil horizons by slope wash and eolian
mantles, and the polygenetic character of many surface remnants preclude the widespread use of soil stratigraphy andmorphology for purposes of correlating surfaces.
Also, we assume that basalt flows that cap flat-lying mesas
flowed down ancestral valley floors, which were at or near thebase level of the ancestral axial stream. Then we assume

that the vertical change in the elevation of the surface (relative to axial-stream base level) is due to tectonism. Problems using basalt flows to locate paleobase-level elevations
and then infer subsequent tectonic deformation include: (1)flows may be of considerably different age between mesas.
Furthermore, age uncertainties for some radiometric dates
cover a large enough time interval for considerable landscape changes to occur, irrespective of tectonism. (2) If theflows occupied tributary canyons, then the height above theaxial-stream base level is still unknown. And (3) initialformation of the surfaces may greatly antedate the flowsthat bury them. Nevertheless, the position of these flows
offers the best proxy for base level and, where dated, themost reliable numerical age constraint for the surfaces theybury. In the absence of better spatial and temporal constraints we will develop our interpretation of drainage andtectonic evolution, fully acknowledging the limitations of
our approach.

Pre-Quaternary surfaces and the Puye
aggradation episode
The oldest geomorphic surface preserved in the study area
is T1. The T1 surface is late Miocene in age and is buriedby the 8 Ma Lobato basalt flows which cap high mesas bothin the Colorado Plateau and on the western edge of theAbiquiu embayment (Fig. 4A, Table 1). The T1 surface
occurs in the Colorado Plateau below the 7.8 ± 0.7 Ma(Manley and Mehnert, 1981; Table 1) Lobato basalt flowthat caps Cerro Pedernal and is underlain by alluvium
which caps the upper Abiquiu Formation (see stratigraphic
column of Church and Hack, 1939). Cerro Pedernal is 9862
ft (3006 m) above sea level and the contact between theburied T1 surface and overlying basalt flow is at
approximately 9700 ft (2960 m) elevation, or, more
importantly, at approximately 3630 ft (1110 m) above thepresent-day Rio Chama (Fig. 4A).
In the Abiquiu embayment, the T1 surface is preserved
beneath basalt flows on parts of Cañons Mesa, Mesa Escoba (7.9 ± 0.5 Ma; Manley and Mehnert, 1981), and thenorthern end of Polvadera Mesa (7.8 ± 0.5 Ma; Manley and
Mehnert, 1981; Figs. 2, 4A). Throughout the Abiquiu embayment, the T1 surface is at approximately 7900 ft (2400m) elevation, or 1760 ft (540 m) above the present-day Rio
Chama (Fig. 4A). We believe the difference in the height of
the Ti surface from the Colorado Plateau to the Abiquiu

embayment is due to tectonism. Evidence for this is presented below.
T2 is an early Pliocene surface found at only one locality
in the study area, near Sierra Negra (Figs. 2, 4B). Sierra
Negra is a volcanic vent at 8052 ft (2454 m) elevation, fromwhich issued a 4.8 ± 0.1 Ma (UAKA-77-93; Baldridge et al.,1980) basalt flow. The basalt flow descends from the ventdown a hillslope towards the southeast, entering an ancestral
stream channel at 7700 ft (2350 m) elevation. Within the
paleochannel, a fluvial gravel deposit occurs beneath theflow (Wilson, 1977; Fig. 4B). The contact between theburied alluvium and the basalt is at approximately

347650 ft (2330 m) elevation, or 1750 ft (530 m) above the
Rio Chama channel (Fig. 4B). Unfortunately, since so littleof the 4.8 Ma tributary channel is preserved and the exact
location of the ancestral Rio Chama is unknown, we cannot
determine the elevation of the axial-stream base level at 4.8
Ma other than to conclude that it must have been lower than
the tributary-valley floor [7650 ft (2330 m) elevation].
T3 is a buried Pliocene surface formed prior to the deposition of the Puye Formation (Figs. 3, 4C). The
importance of identifying the buried T3 surface is threefold: first, to call attention to a Pliocene aggradational
episode; second, to support an argument that the Embudo
and Pajarito fault zones were actively moving (with downto-the-south and -east motion, respectively) during the
Pliocene; and third, to illustrate the role of intra-basin
tectonism on the formation and preservation of geomorphic
features. It has been recognized only south of the Embudo
fault zone, and may be entirely absent in the Abiquiu
embayment and on the Colorado Plateau where we havefound no evidence of Pliocene aggradation.
Pliocene aggradation is recorded by the formation oflarge fan complexes on both the east and west sides of theEspañola basin. On the east side the Ancha Formationconsists of sediment derived from the Sangre de Cristo
Mountains. A Cerros del Rio basalt flow (1.96 ± 0.06 Ma;
Manley, 1976b) that covers part of the upper Ancha
Formation constrains fan formation to the Pliocene. The
Puye Formation was deposited during the Pliocene on thewest side of the Española basin (Fig. 4C). The Puye
Formation is a volcano-genic, rift-margin fan complex
derived from the Jemez Mountains volcanic pile primarily
during and after the deposition of the Tschicoma Formation(7-3 Ma; Gardner and Goff, 1984; Waresback, 1986;
Waresback and Turbeville, 1990). The Totavi Lentil (amember of the Puye Formation named by Griggs, 1964)
consists of mid-Pliocene Rio Grande (or Chama) axialstream deposits. The Puye Formation interfingers with
Cerros del Rio basalt flows (Fig. 4C). Radiometric ages
from the Cerros del Rio basalt provide temporal constraints
on the timing of aggradation of the Puye and Ancha fans.
Between the Embudo fault zone and White Rock Canyon,
the base of the Puye Formation (i.e. the buried T3 surface) isin places about 100 ft (30 m) above the Rio Grande (e.g.
Ancho Canyon; Dethier, in press; Fig. 4C). Just 2-3 mi (35km) upstream along some of the tributaries (e.g. Guaje
Canyon, Chupaderos Canyon, Garcia Canyon, etc.), the baseof the Puye Formation lies below the level of tributarystream valley floors, either because the basal Puye is backtilted by down-to-the-east motion on the Pajarito fault zone
or because the Puye fan was deposited on a lower-gradientsurface than is that of modern canyons draining the Pajarito
Plateau (Griggs, 1964; also see cross-section B-B' of Smithet al. , 1970; Fig. 4C).
In contrast, the Puye Formation (and the upper Santa Fe
Group) is considerably higher throughout the Abiquiu embayment north of the Embudo fault zone (contrast crosssection A-A' with B-B' of Bailey et al., 1970). Also, thePuye Formation is thin north of the Embudo fault zone [lessthan 100 ft (30 m) thick as compared to 1000 ft (300 m) or
more in places south of the Embudo fault]. Furthermore,
only isolated remnants of the Puye Formation are foundnorth of the Embudo fault zone, suggesting that the locus ofdeposition lay to the south, and little aggradation occurred in
the Abiquiu embayment during the Pliocene. The striking

contrast between large-scale fan aggradation south of thefault (see Fig. 4C) and general base-level stability
northwest of the Embudo fault zone (see Fig. 4B) during
the Pliocene may suggest tectonic activity on the Embudo
and Pajarito faults beginning in the Pliocene.
The T4 surface may mark an extended period of
Pliocene base-level stability in the Abiquiu embayment.
The T4 surface has been preserved by 3.2-2.8 Ma El Altobasalt flows on Cañones Mesa and Mesa de Abiquiu, andthe 2.8 Ma Servilleta basalt flows that cap Black Mesa (Fig.
4B). We believe these basalt flows coursed down Pliocene
stream valleys at or near an ancestral base level of the Rio
Chama/ Rio Grande for the following reasons.
(1) All the flows have low surface gradients which decline toward, or parallel to, the axial drainage. The flow
capping Mesa de Abiquiu (3.2 ± 0.1; UAKA-77-85; Baldridge et al., 1980) has a gradient of 2.2% in its upper reach
before decreasing to 0.7% in its terminal reach. The Black
Mesa flow has a surface gradient of approximately 1.6% in
its southern end, which presumably followed the ancestral
Rio Grande channel.
(2) Well-rounded stream gravel (in-situ and in float) ispresent below several of the flows. For example, Manley(1976a, pp. 43-45) reports finding axial-stream gravel beneaththe Servilleta basalt flow (2.78 ± 0.44; Manley, 1976b) thatcaps Black Mesa. She states that the "composition,
imbrication, and rounding of the gravels beneath the Servilleta
Fm. on Black Mesa indicate the presence of an axial river withsource areas in the Sangre de Cristo Mountains and north of
the Española basin" (Manley, 1976a, p. 45). Also, we havenoted a considerable quantity of well-rounded cobbles in thescree accumulated at and below the base of the basalt flow on
Cañones Mesa. The clast lithologies are diagnostic of Rio
Chama axial-stream deposits (i.e. Precambrian quartzites,
crossbedded quartzites, and quartzitic conglomerates). Since
no axial-stream deposits rest above the El Alto basalt flow, we
argue that the flow capping lower Cañones Mesa has buried aformer channel of the Rio Chama.
(3) The width of the flows is comparable to the width ofthe present-day drainages.
Quaternary surfaces and periods of aggradationand incision
Pre-Bandelier valley incision is recorded primarily along
the northern and eastern flanks of the Jemez Mountains.
Incision began after deposition of the Puye Formation[sometime after 1.75 Ma, based on the age of a rhyolite
plinian deposit in the uppermost part of the Puye Formation(1.75 ± 0.08 Ma; Spell, unpublished data reported by Waresback and Turbeville, 1990)]. Incision ended before 1.50 Ma (± 0.01 Ma; Spell et al., 1989), when lower Bandelier Tuff
sediment began to accumulate on the floors of existing canyons (Fig. 4D). During this interval of approximately
200,000 years, streams draining the eastern and northernflanks of the Jemez Mountains cut canyons which in places
were as deep or deeper than present-day canyons (Dransfieldand Gardner, 1986; Fig. 4D).
The pre-Bandelier canyons were quickly back-filled with
Bandelier pyroclastic ejecta from 1.5 to 1.13 ( ± 0.01) Ma(Spell et al., 1989; Fig. 4D). During this aggradation, thedeeply dissected pre-Bandelier landscape was quickly transformed to a low-relief, gently sloping (2-3°) surface (Dransfield and Gardner, 1985). Furthermore, minor incision may
have also occurred during general aggradation between dep-

35
osition of the lower and upper members of the Bandelier
Tuff as the drainage systems of the Pajarito Plateau began
to re-establish themselves.
A mid- to late Quaternary period of incision followed theperiod of Bandelier aggradation and resulted in re-excavation of canyons and deposits in the Abiquiu embayment andEspañola basin (Fig. 4E). Post-Bandelier incision appears
to have continued since 1.1 Ma except for minor base-levelstands or aggradational periods that were sufficiently longto allow Quaternary surfaces (on pediments, strath and fillterraces) to form.
Ten Quaternary-age surfaces are recognized in the
study area (Fig. 4E). Seven Quaternary surfaces appear
with regularity throughout the Colorado Plateau andAbiquiu embayment. We tentatively label these Q1–Q7,
oldest to youngest. In addition, there are three Quaternarysurfaces that predate Q1 (here tentatively labeled Qg,
Qoal, and undesignated) (Fig. 4C, D). We describe andsummarize the ages (where possible) of the Quaternarysurfaces in Table 2.
Attempts to correlate the surfaces throughout the study
area are difficult due to the spatial variation in the elevation
at which these surfaces occur and due to the limited utilityof other correlative tools (e.g. soil morphology). Variationof the number of surfaces and their height above modernday river grade may result from:(1) Continued tectonic movement on faults following
formation of the surfaces.

(2) Changes in the modem-stream and paleostream gradients due to lithologic control on stream incision.
(3) Variability due to polygenetic surface evolution. For
example, many surface remnants north of the Rio Chama in
the Plaza Colorado and Plaza Blanca land grants (Stop 8)
are topographically isolated buttes which now experience nosurface runoff from adjacent highlands. In contrast, many
surfaces south of the Rio Chama are contiguous with theflanks of the Jemez Mountains, thereby receiving runoff
from the adjacent highlands. Eolian processes may be relatively more important on the surfaces located on buttes,
while overland flow may be an important surface-modifying
process on surfaces which receive runoff from adjacenthighlands.
(4) The possibility that surfaces diverge or converge ineither the downstream or upstream direction due to changes
in sediment load and transport ability (Harvey et al., 1988).
(5) Measurement errors and investigator biases—one of
us differentiating surfaces more finely than the other.
The sedimentary deposit associated with unit Qoal is a
stratigraphic enigma. It consists of Rio Grande axial gravel
found on top of a 1.8 Ma (Dethier, submitted; Table 1)
Cerros del Rio basalt flow near the White Rock Canyon
overlook (Stop 9; Fig. 4C). The perplexing question iswhether this gravel was deposited near the end of Puye
aggradation and prior to pre-Bandelier valley incision, in
which case it could be considered a part of the Puye
Formation, or in the mid-Quaternary following Bandelier
aggradation (1.5-1.1 Ma; Spell et al., 1989). Mineralogic
and/or chemical analyses of the interstitial material within
the deposit may resolve the question of age.
Deposits of unit Qg comprise a single exposure of gravel
which is buried by the Guaje pumice (Fig. 4D), the basal
bed of the Bandelier Tuff, indicating the gravel is at least
1.5 Ma old (Table 2). This gravel has been previously recognized by Hawley (1978). It demarcates a stream-channel
position before Bandelier aggradation commenced 1.5 Ma.
Unit Q1 is a high erosion surface cut on the Bandelier
Tuff along the northern and eastern flanks of the Jemez
Mountains. In the Colorado Plateau and in parts of theAbiquiu embayment the Q1 unit forms a terrace strath underlain by an axial-gravel deposit 1-5 m thick. The age of theQ1 surface is less than 1.1 Ma and greater than 0.62 Ma
based on stratigraphic relationships with the Bandelier Tuffand Lava Creek B ash (Table 2; Fig. 4D, E). In the Colorado
Plateau, the height of the Q1 surface is. usually 465 ± 30 ft(140 ± 10 m) above stream grade. Further downstream
within the western half of the Abiquiu embayment, the Q1
surface is usually 430 ft (130 m) above grade. Downstream
from the Embudo fault zone, Dethier et al. (1988) mapped
the Q1 surface at approximately 500 ft (150 m) above thepresent-day stream gradient. Much of this variation insurface elevation may be due to lithologic variations between
the Colorado Plateau and the Rio Grande rift, and the river's
ability to cut through the different lithologies. Also, there isan elevation uncertainty whenever projecting remnants of a
sloping surface. In addition, Quaternary fault movements
could be responsible for variation in surface heights (asappears to be the case at Stop 8).
Unit Q2 is often recognized in the field by the presence ofreworked Lava Creek B ash (620 ka) just above the axialstream gravel deposits (Figs. 4E, 5; Dethier et al., 1988). The
ash provides a valuable mid-Quaternary isochron for regionwide correlation (Table 3). As yet, our investigations

have failed to locate any ash deposits along the margin of
the Colorado Plateau. Nevertheless, widespread occurrenceof surface remnants at a consistent height above the presentday axial stream allows the Q2 surface to be traced fairly
confidently from the rift onto the Colorado Plateau. In the
Colorado Plateau, the Q2 surface lies 405 ± 25 ft (120 ± 10m) above stream grade. In the Abiquiu embayment, the Q2
surface is generally 350 ± 25 ft (110 m) above stream grade;and near the Embudo fault, the Q2 surface is about 410360
ft (120-110 m) above grade (Table 2, Fig. 4E).
The remaining Quaternary surfaces are formed on dep

ositional insets at elevations between the Q2 surface andthe modern channel (Fig. 6). We believe each successive
surface is younger than the surfaces above. Table 2 summarizes the data that limit the ages of these mid- to late
Pleistocene surfaces.
Climate change and development of surfaces
Geology at several stops on this field trip demonstrates
that the Rio Chama has been cutting down since pre-Bandelier time and that incision has been interrupted periodically by intervals of relatively stable base level. Uplift ofboth the Colorado Plateau and the Rio Grande rift may havehelped to drive incision by the Rio Chama. However, climate change must have also contributed to downcutting andhas likely mediated periods of rapid base-level change. The

37change from relatively equitable climate to alternating glacial
and interglacial periods at about 2.0 Ma probably increased
the total flow in the Rio Chama system and the magnitude andduration of peak flows during the snowmelt period. Climatechange also would have affected vegetation cover, sedimentavailability, and bed material, variables that help to control
the balance between aggradation, lateral planation, andincision. We have not collected data that give us direct
information about these variables; however, we hypothesizethat much of the incision recorded along the Rio Chama is a
result of higher peak discharge since mid-Pleistocene time.
As the Rio Chama incised through Miocene bedrock, at leastnine similar stratigraphic sequences were deposited atelevations ranging from 15 to 170 m above

present grade (Fig. 6). A flight of these surfaces is wellpreserved in three unnamed arroyos located south of thejunction of US-84 and NM-554 (NM-96 on older maps; Fig.
7). Stop 3 traverses the Quaternary surfaces, which are
underlain by axial-stream and valley-margin fan deposits. The
typical sedimentary sequence of deposits underlying mid- tolate Quaternary surfaces comprises: (1) a basal cobble gravel
deposited by the ancestral Rio Chama on a bedrock surface;(2) 2-4 m of silty sand and sand deposited as overbank andalluvial-fan deposits; and (3) 5-40 m of medium sand to cobble
gravel deposited by alluvial fans (Fig. 5). This typicalsedimentary sequence is best preserved within the Rio Grande
rift, south and west of the Rio Chama, where the tributaries
draining the Jemez Mountains provided the38coarse alluvial-fan materials that cap and preserve the finegrained overbank deposits. North of the Rio Chama andthroughout most of the Colorado Plateau margin, the sedimentary sequence usually consists of the basal cobble
axial-stream gravel with an eolian mantle of 0-2 m thickpreserved in places.
Lava Creek B ash and, at many localities, fossil gastropods are preserved in the middle layer of this sequence (Fig.
5). Field relations suggest that alluvial fans prograded acrossthe ancestral Rio Chama floodplain during episodes when
the Rio Chama meander belt narrowed and shifted to thenorth and east. Subsequent lateral planation and downcuttingby the ancestral Rio Chama broadened the meander belt anddeepened the valley before alluvial fans again advanced
across the valley. Love et al. (1987) used mapping andsubsurface evidence to demonstrate that near Española,alluvial-fan progradation and narrowing of the meander belt
began in the latest Pleistocene (at approximately 15 ka) andmay have continued through the early Holocene. If thispattern of deposition and those described by Gile et al.
(1981) in the Desert Project area of southern New Mexico
are reasonable analogs for the Chama valley, then the repeated, inset sequences of ancestral Rio Chama sedimentprobably are related to episodic climate change. As a working hypothesis, we suggest that fan construction occurs during the transition from cool, moist to warm, dry climate, thataxial-river incision occurs during the early or middle part ofglacial pluvial cycles, and that minor aggradation andincision or lateral planation dominate most of the time.
We can use the deposits exposed at Stop 1 (near the
Española High School) and two radiocarbon dates to constrain the timing of incision during the most recent period ofclimate change. At about 26 ka the Rio Chama flowed about
15 m above what Love et al. (1987) mapped as the latest
Pleistocene level. Amino-acid ratios in higher elevation
deposits (D. P. Dethier, unpublished data 1988) suggest thatsince the early Wisconsinan the Rio Chama had flowed
within 10 m of its elevation at 26 ka. An apparently brief
period of alluvial-fan aggradation occurred after 26 ka, andwe interpret dune activity at 19 ka (Stop 1; Fig. 8) as
evidence that the Rio Chama meander belt had widened
following the downcutting that occurred between 26 and 19

ka. Strong eolian activity marked the transition from a pluvial to a more arid climate (see also Wells et al., 1990) that
resulted in alluvial-fan construction from after 19 ka to theearly Holocene(?). We infer from these data that incisionby the axial channel most likely occurred close to the beginning of the pluvial period.
Neotectonic analysis of faults of theAbiquiu embayment
The Abiquiu embayment is a geophysically distinct segment of the Rio Grande rift (Biehler et al., in press), separated
from the Colorado Plateau by NNE-trending normal faultsand from the remainder of the Española basin by the ENEtrending Embudo fault zone. The Embudo fault zone appears
to be a transfer fault as defined by Gibbs (1984; alternatively
referred to as an intracontinental transform fault byMuehlberger, 1979). The Embudo fault zone accommodates
differences in crustal extension on each side of the fault(Aldrich, 1986) and transfers motion between rift grabens(Española and San Luis basins). The Abiquiu embayment,
northwest of the Embudo fault zone, is shallower than the restof the Española basin. The thickness of Tertiary rift-fillsediment and Paleozoic/Mesozoic sedimentary rocks above
the Precambrian basement is approximately 1 km (3000 ft;
unpublished seismic profiles produced by SAGE in the summer of 1990; W. S. Baldridge, pers. comm. 1990). In contrast, the Española basin southeast of the Embudo fault zone
has 2-3 km (7000-10,000 ft) of strata above the Precambrianbasement (Biehler et al., in press).
Seven fault zones along the margin of the Rio Grande rift
and within the Abiquiu embayment have been examined. The
objectives are: (1) to determine the geometry of the fault
zones; and (2) to find kinematic indicators along fault zones.
The purposes are to determine the timing, magnitude, and rateof faulting along this rift margin and then to relate the local
features to the regional tectonic forces affecting the evolutionof the central Rio Grande rift in northern New Mexico. Thesefaults within the Abiquiu embayment are informally named(from west to east) the Cañones, Garcia, Cerrito de la
Ventana, Cerrito Blanco, Plaza Colorado, Madera Cañon, andEmbudo fault zones (see Fig. 2 for locations). Over 450 faultplane attitudes (main, antithetic, and cross faults; R1 and R2
planes to main, antithetic, and cross faults) and 300 faultplane lineations (slickensides, striations, oriented mineral
growths, and aligned particles) have been collected. The
structural data for six of the fault zones, where data aresufficient to define trends, are depicted graphically instereonets (Fig. 9A—L) and tabulated numerically in Table 4.
In the following paragraphs, we report preliminary results of
our ongoing fault-plane analyses.
Cannes fault zone
The Cannes fault zone (mapped by Manley, 1982) is thewesternmost rift-bounding structure in the Abiquiu embayment. We informally divide the Cannes fault zone into awest splay (Gonzales fault of Kelley, 1979), east splay(Cañones fault of Kelley, 1979) and main splay (Cobre
fault? of Kelley, 1979). The main splay of the Cañones fault
zone crosses the Rio Chama valley and US-84
approximately 5 mi (8 km) west of Abiquiu (Stop 6; Fig. 2).
At this locality the Cañones fault zone has at least 400 ft(120 m) of strati-graphic offset, juxtaposing upper Paleozoic
and Mesozoic strata (Cutler and Chinle Formations) on thewest against Eocene to early Miocene rift-fill sediments (ElRito andlower Abiquiu Formations; Fig. 3) on the east. This estimateof stratigraphic throw at the surface is in good agreementwith subsurface seismic imaging of the fault done by the
1990 SAGE field class (W. S. Baldridge, pers. comm. 1990,
and unpublished seismic data).
The average rake of lineations on the main fault of the
Cañons fault zone is 79° NNE, indicating that the fault zone
has formed primarily through dip-slip movement, alongwith a small component of left-lateral slip on the main fault.
[Note: average rake angle is always acute. Rake angle =(tan of angle between fault-plane strike and lineation trend)/(cos of the dip angle of the fault plane).]Vertical motion along the Cañones fault zone appears toincrease to the south where the fault zone bifurcates into thewest and east splays. We use the deformation of the T1
surface preserved beneath the late Miocene basalt flows on
Cerro Pedernal, Mesa Escoba, and Polvadera Mesa to estimate that there is approximately 1870 ft (570 m) ofvertical offset across the east and west splays of the Cañons
fault zone (Figs. 2, 4A). The amount of throw is calculated
relative to the present-day Rio Chama gradient, which we
assume is comparable to the late Miocene gradient of the

Rio Chama. The long-term vertical slip rate across the Cañones fault zone is approximately 0.007-0.008 cm/yr (600 m/ 8my). This long-term rate is a minimum slip rate since wecannot constrain when after 8 Ma motion began and whenmotion ceased (if it has) along this fault zone.
Sedimentary units, traditionally thought of as rift-fill sediment (i.e. the Eocene El Rio Formation and late Oligocene toearly Miocene Abiquiu Tuff of Smith, 1938; Fig. 3) are foundboth east and west of the Cannes fault zone. Interpretation ofthis spatial pattern by previous workers suggested that theearly rift-filling sedimentation (El Rito and Abiquiu
Formations of Smith, 1938) took place on the broadly warped
eastern limb of the Nacimiento uplift (a Laramide structureon the eastern edge of the Colorado Plateau) and the west
margin of the Española basin before late Miocene andPliocene faulting created a morphologically distinct rift basin(May, 1980, 1984; Baltz, 1978; summary of Wood andNorthrop, 1946, and Smith et al., 1970).
Garcia and Cerrito Blanco fault zones
The Garcia and Cerrito Blanco fault zones lie several
kilometers east of the Cañones fault zone. Both fault zones

42
trend NNE (N 15-27° E) with normal faults dipping 6470°
to the ENE. Antithetic faults, located tens to hundreds ofmeters east and parallel to the main fault, form graben
structures along the faults. In addition, the two fault zones
are connected by Cerrito de la Ventana, a basaltic dike andfault zone which trends approximately N 35-45° E. The
orientation of this dike and fault may have kinematic significance. The Garcia fault offsets upper Abiquiu Formation
south of the Rio Chama and juxtaposes lower Abiquiu Formation (on the west side of the fault) against upper Abiquiu
Formation (on the east side). The Cerrito Blanco fault juxtaposes upper Abiquiu Formation (west side) against the Ojo
Caliente Member of the Tesuque Formation (east side)along Cerrito Blanco. North of Cerrito Blanco, near theintersection of Cerrito Blanco fault zone with Cerrito de la
Ventana, lower Abiquiu Formation and El Rito Formation(on the west side) are juxtaposed against the Ojo Caliente
and ChamaEl Rito Members of the Tesuque Formation andthe upper Abiquiu Formation. In general, the farther northalong this fault and along Cerrito de la Ventana, the older
are the strata offset by the faults.
Lineations on the fault planes of the Garcia fault zone
indicate that dip-slip movement is dominant, but with aminor component of right-lateral slip (average rake of lineations on fault plane is 84° SSW). Lineations on fault planesof the Cerrito Blanco fault zone also indicate dominant dipslip movement, but with a minor left-lateral slip (averagerake of lineations is 83° SSE). Normal faulting on these two
fault zones appears to have thrown the T4 surface (preservedon Cañones Mesa and Mesa de Abiquiu by burial with 3.22.8 Ma El Alto basalt flows) 560 ft (170 m) relative to thepresent-day Rio Chama surface. Combined the longterm
vertical slip rate for the Garcia and Cerrito Blanco fault zones
is 0.006 cm/yr (170 m/3 my). Individually, the Garcia fault
zone appears to have displaced a Lobato basalt flow [3 mi (5km) southwest of Abiquiu] by approximately 250 ft (80 m),
suggesting that the average vertical slip rate for the Garcia
and Cerrito Blanco fault zones individually is about 0.003
cm/yr (80-90 m/3 my). As these faults do not appear todisplace Quaternary terraces in the Rio Chama valley, theselong-term slip rates are minimum values.
Plaza Colorado fault zone
The Plaza Colorado fault zone strikes N 75° E and dips
68° SSE, juxtaposing upper Abiquiu Formation on both sidesof the fault zone. Lineations on the fault-plane surface havean average rake of 77° ENE, again suggesting that movementon the fault is primarily dip-slip, with a minor component ofleft-lateral slip. We have been unable to calculate a slip rate
for the Plaza Colorado fault zone because we have not
located a piercing point nor observed a displacement of welldeveloped, dated surface.
Madera Cañon fault zone
The Madera Cañon fault zone strikes NE. The main fault
plane dips an average of 62° SE. An antithetic fault planestrikes nearly parallel to the main fault trace and dips NW.
The result is the formation of a distinct graben within thefault zone (Stop 8). The fault zone juxtaposes upper Abiquiu
Formation on the northwest side of the fault zone against the
Chama—El Rito Member of Tesuque Formation on thesoutheast side. The average rake of fault-plane lineations is
86° NE, indicating that the graben formed from nearly pureextension with only a minor amount of left-lateral slip.

The Madera Cañon fault zone exhibits late Pleistocene
movement with disruption of the unit Q4 (Pinedale-equivalent surface?) but not the immediately younger unit Q5(Fig. 10; Stop 8). The Q4 pediment surface on the west sideof Madera Cañon has been offset about 30 ft (10 m), while
the longitudinal profile of unit Q5 is undeformed, indicatingthat the last detectable movement on the Madera Cañon
fault zone postdates formation of unit Q4, but antedates
formation of the Q5 surface. A tentative late Pleistocene
slip rate on this fault is about 0.01-0.007 cm/yr (10 m/
approximately 80-150 ka; see Table 2 for age estimates of
Quaternary surfaces).
Embudo fault zone
The Embudo fault zone generally trends ENE and dips an
average of 75° SSE. In the vicinity of Arroyo de la Presa(Stop 2), a series of NNE-trending cross-fault segments
connects left-stepping, en-echelon, ENE-trending main-fault
segments. Miocene sedimentary rocks and basalt flows andPliocene(?) rocks south of the zone dip steeply to the south
and are overturned locally. The Ojo Caliente Member of
Tesuque Formation and the lower Chamita Formation cropout north of the fault (Dethier and Manley, 1985; Aldrichand Dethier, 1990), but are not exposed to the south.
The Embudo fault zone differs from the other faults inseveral important aspects. First, based on seismic images
and gravity measurements, it appears to have a significant
structural relief (second only perhaps to the Cañones faultzone). Second, right-lateral rather than left-lateral slip characterizes movement on main-fault segments in the Arroyo
de la Presa area (Aldrich, 1986; Stop 2). However, theaverage rake of lineations on the short, NNE-trending crossfault segments is 32° NNE and yields a left-lateral sense of
movement. Unfortunately, we have not yet found sufficient
fault-plane data to rigorously quantify movement on the
ENE-trending main-fault segments of the Embudo fault
zone. Finally, the Embudo fault zone appears to be an
active transfer fault, accommodating differences between
the faster spreading rates in the Española basin and theslower spreading rates in the Abiquiu embayment and San
Luis basin (Aldrich, 1986).
Objectives of our ongoing structural analyses include
enhancement of existing fault-plane data with more measurements from key fault localities which are currently under-represented in our data set (e.g. the ENE-trending mainfault segments of the Embudo fault zone). In addition,
microscopic examination of oriented fault-gouge samples

43
have begun, along with observations of megascopic features, such as the attitude of dikes which presumably are
oriented in a plane orthogonal to the minimum
compressive stress(?).
With enhanced data sets of micro-, meso-, and megascale
kinematic indicators, the regional tectonic significance of thelocal structures can be addressed. Important questions toaddress include the temporal significance of structures andthe state of the regional stress field affecting the central Rio
Grande rift of northern New Mexico during the Neogene andQuaternary periods. Specifically, did NNE-trending faults
form coevally or successively, as suggested by the offset ofprogressively younger strata in an easterly direction? Are the
NNE- and ENE-trending fault sets conjugate, formed in
response to a single, mutually experienced stress field or,alternatively, in response to progressive change in the regional stress field (perhaps due to the 45° clockwise rotationof the Colorado Plateau reported by Zoback and Thompson,
1978; Zoback and Zoback, 1980; Zoback et al., 1981; Aldrich and Laughlin, 1984; Aldrich, 1986)? Does the fault
pattern reflect transtension with formation of a pull-apartbasin? Finally, is the megascale pattern congruous with an
intra-rift transfer structure (i.e. Embudo fault zone) that
accommodates en-echelon offset of rift grabens and differences in extension on either side of the accommodation zone
(e.g. Aldrich, 1986)?
Conclusions and future work
The geomorphic history of the Abiquiu embayment andeast flank of the Jemez Mountains comprises two periods ofaggradation (the Puye and Bandelier aggradations); three
periods of incision (the late Miocene through Pliocene, preand post-Bandelier periods of incision); four pre-Quaternary
geomorphic surfaces (T1—T4, designated from oldest to
youngest) and as many as 10 Quaternary surfaces. These
geomorphic features record multiple cycles of large-scaleaggradation and incision along the northern and eastern
flanks of the Jemez Mountains, suggesting that theconstruction of the volcanic pile and the voluminous eruptionof volcanic material have had an impact on thegeomorphology of the area.
Structural development of a distinct morpho-tectonic riftbasin began in the late Miocene (Baltz, 1978; Wood andNorthrop, 1946) with the onset of Lobato volcanism and thepostulated concomitant movement on the Cañones fault
zone. We infer from the apparent displacement of the T4
surface that subsequent movement occurred after 3 Ma on
the Garcia and Cerrito Blanco fault zones. Finally, the Embudo and Pajarito fault zones seem to have been active
during the Pliocene and Quaternary (approximately 4-5 Ma;Manley, 1976a, 1978; Golombek, 1983; Gardner and Goff,
1984; Aldrich, 1986). The Embudo fault zone appears to bea transfer fault which segments the Española basin into theshallow Abiquiu embayment to the northwest and a deeper
basin to the southeast. This intra-basin tectonism may haveimportant consequences on the initial formation,subsequent preservation, and eventual correlation of geomorphic and sedimentologic features between basins aswell as within a single basin.
The significance of two distinct fault sets (one oriented
roughly NNE—SSW and the other ENE—WSW) is still
enigmatic. Specifically, do these fault sets reflect: (1) acoeval conjugate set of related faults formed in response to
one stress field; (2) a coeval set of related faults formed fromlateral-slip (or oblique-slip) motion within or adjacent to apull-apart basin; or (3) a time-transgressive set of faults
that formed as the regional stress field changed due toclockwise rotation of the Colorado Plateau.
Better temporal control on mid- to late Quaternary surfaces would assist in region-wide correlation of surfaces,
better address roles of climatic—tectonic—volcanic
perturbations in landscape evolution, and allow for
calculation of rates of landscape evolution. Such time
control is currently being pursued with attempts to use Useries disequilibria dating of carbonate rinds formed on
clasts in terrace deposits.
Acknowledgments
We would like to thank S. G. Wells, C. Mawer (University of New Mexico), C. D. Harrington, W. S. Baldridge,
J. N. Gardner, M. J. Aldrich, Jr., S. L. Reneau (Los Alamos
National Laboratory), and J. W. Hawley (New Mexico Bureau of Mines & Mineral Resources) for advice, comments,
and ideas shared with us during our years of work in the
Rio Grande rift. We are grateful to S. G. Wells and C. D.
Harrington for their critical reviews of an earlier draft ofthis paper.
Grant Meyer (University of New Mexico) provided Gonzalez with invaluable field assistance and entertained him
with his humor. Anthony Garcia and Eric Montoya (Los
Alamos National Laboratory) drafted figures and maps, andBarbara Hahn and Josephine Ortega (Los Alamos National
Laboratory) prepared the tables. We are most grateful for
their professional assistance and patience.
Financial support to Gonzalez for field work came from
Sigma Xi, the Scientific Society, New Mexico GeologicalSociety, Geological Society of America, and the Geology
Department and the Student Research Allocations Fund of
the University of New Mexico. The Earth and Environmental Sciences Division of the Los Alamos National Laboratory provided invaluable logistical support for field
operations and equipment for laboratory analyses.
References
Aldrich, M. J., Jr., 1986, Tectonics of the Jemez Lineament in the Jemez
Mountains and Rio Grande rift: Journal of Geophysical Research,
91(B2): 1753-1762.
Aldrich, M. J., Jr., and Laughlin, A. W., 1984, A model for the tectonic
development of the southeastern Colorado Plateau boundary: Journal ofGeophysical Research, 89(B 12): 10,217-10,218.
Aldrich, M. J., Jr., and Dethier, D. P., 1990, Stratigraphic and tectonic
evolution of the northern Espanola basin, Rio Grande rift, New
Mexico: Geological Society of America, Bulletin, 102: 1695-1705.
Aldrich, M. J., Jr., Laughlin, A. W., Meade, J. S., and Peirce, H. W.,
1986, The Jemez Lineament: Structural boundaries and control on sedimentary facies, tectonism, and mineralization: Proceedings of the 6thInternational Conference on Basement Tectonics, pp. 104-113.
Bachman, G. 0., and Mehnert, H. H., 1978, New K—Ar dates and the late
Pliocene to Holocene geomorphic history of the central Rio Grande
region, New Mexico: Geological Society of America, Bulletin, 89: 283292.
Baldridge, W. S., Bartov, Y., and Kron, A., 1983, Geologic map of the Rio
Grande rift and southeastern Colorado Plateau, New Mexico andArizona; supplement to Riecker, R. E., (ed.), Rio Grande rift: tectonics
and magmatism: American Geophysical Union, Washington, D.C.
Baldridge, W. S., Olsen, K. H., and Callender, J. F., 1984, Rio Grande
rift: Problems and perspectives: New Mexico Geological Society,
Guidebook 35: 1-12.
Baldridge, W. S., Damon, P. E., Shafiqullah, M., and Bridwell, R. J.,
1980, Evolution of the central Rio Grande rift, New Mexico: New
potassium—argon ages: Earth and Planetary Science Letters, 51: 309321.
Baltz, E. H., 1978, Resume of Rio Grande depression in north-central

44
New Mexico: New Mexico Bureau of Mines & Mineral Resources,
Circular 163: 210-228.
Biehler, S., Ferguson, J., Baldridge, W. S., Jiracek, G. R., Aldern, J. L.,
Martinez, M., Fernandez, R., Romo, J., Gilpin, B., Braile,
L. W., Hersey, D. R., Luyendyk, B. P., and Aiken, C. L. (in press), A
geophysical model of the Espanola basin, Rio Grande rift, New Mexico: Geophysics.
Brown, L. L., and Golombek, M. P., 1986, Block rotations in the Rio
Grande rift, New Mexico: Tectonics, 5: 423-438.
Bryan, K., 1938, Geology and ground-water conditions of the Rio Grandedepression in Colorado and New Mexico; in U.S. National Resources
Planning Board, The Rio Grande Joint Investigation in the Upper Rio
Grande basin: U.S. Government Printing Office, 1(2): 197-225.
Bryan, K., and McCann, F. T., 1936, Successive pediments and terracesof the upper Rio Puerco in New Mexico: Journal of Geology, 44: 145172.
Bryan, K., and McCann, F. T., 1938, The Ceja del Rio Puerco, a borderfeature of the Basin and Range province in New Mexico: Pt. 2, Geomorphology: Journal of Geology, 46(1): 1-16.
Cabot, E. C., 1938, Fault border of the Sangre de Cristo Mountains northof Santa Fe, New Mexico: Journal of Geology, 46: 88-105.
Church, R. S., and Hack, J. T., 1939, An exhumed erosion surface in theJemez Mountains, New Mexico: Journal of Geology, 47: 613-629.
Dethier, D. P. (submitted), Geologic map of the White Rock quadrangle:
New Mexico Bureau of Mines & Mineral Resources, scale 1:24,000.
Dethier, D. P., and Demsey, K. A., 1984, Erosional history and soil
development on Quaternary surfaces, northwest Espanola basin, New
Mexico: New Mexico Geological Society, Guidebook 35: 227-233.
Dethier, D. P., and Harrington, C. D., 1987, Geologic dating of lateCenozoic erosion surfaces near Espanola, New Mexico; in Menges,
C. M. (ed.), Quaternary tectonics, landform evolution, soil chronologies
and glacial deposits—northern Rio Grande rift of New Mexico: Friendsof the Pleistocene Field-trip Guidebook, pp. 111-131.
Dethier, D. P., and Manley, K., 1985, Geologic map of the Chili quadrangle, Rio Arriba County, New Mexico: U.S. Geological Survey, Miscellaneous Field Studies Map MF-l814, scale 1:24,000.
Dethier, D. P., and Martin, B. A., 1984, Geology and structure along partof the northeast Jemez Mountains, New Mexico: New Mexico Geological Society, Guidebook 35: 145-150.
Dethier, D. P., Harrington, C. D., and Aldrich, M. J., 1988, Cenozoic
erosion rates, Espanola basin, New Mexico: Geological Society ofAmerica, Bulletin, 100: 928-937.
Dethier, D. P., Halverson, N., Marrack, L., Meagher, M.. Oelkers, E.,
Harrington, C. D., Sarna-Wojcicki, A., and Meyer, C. E., 1990, Occurrence of the Lava Creek B tephra layer in the northwestern Espanolabasin, New Mexico: New Mexico Geology, 12(4): 77-82.
Dransfield, B. J., and Gardner, J. N., 1985, Subsurface geology of the
Pajarito Plateau, Espanola basin, New Mexico: Los Alamos National
Laboratory, Report LA-10455-MS.
Galusha, T., and Blick, J. C., 1971, Stratigraphy of the Santa Fe Group,
New Mexico: American Museum of Natural History, Bulletin, 144: 1127.
Gardner, J. N., and Goff, F., 1984, Potassium–argon dates from the Jemez
volcanic field: Implications for tectonic activity in the north-central Rio
Grande rift: New Mexico Geological Society, Guidebook 35: 75-82.
Gardner, J. N., and House, L., 1987, Seismic hazards investigations at
Los Alamos National Laboratory, 1984-1985: Los Alamos National
Laboratory, Report LA-1 1072-MS.
Gibbs, A. D., 1984, Structural evolution of extensional basin margins:
Geological Society of London, Journal, 141: 609-620.
Gile, L. H., Hawley, J. W., and Grossman, R. B., 1981, Soils and geomorphology in the Basin and Range area of southern New Mexico—
Guidebook to the Desert Project: New Mexico Bureau of Mines &Mineral Resources, Memoir 39, 222 pp.
Golombek, M. P., 1983, Geology, structure, and tectonics of the Pajarito
fault zone in the Espanola basin of the Rio Grande rift, New Mexico:
Geological Society of America, Bulletin, 94: 192-205.
Golombek, M. P., McGill, G. E., and Brown, L., 1983, Tectonic andgeologic evolution of the Española basin, Rio Grande rift: Structure,
rate of extension, and relation to the state of stress in the western
United States: Tectonophysics, 94: 483-507.
Griggs, R. L., 1964, Geology and ground-water resources of the Los
Alamos area, New Mexico: U.S. Geological Survey, Water-Supply Paper 1753, 107 pp.
Harrington, C. D., and Aldrich, M. J., 1984, Development and deformation of Quaternary surfaces on the northeastern flank of the Jemez
Mountains: New Mexico Geological Society, Guidebook 35: 235-239.
Harvey, M. D., Germanoski, D., and Pitlick, J., 1988, Terrace-forming

processes in modern fluvial systems: Implications for Quaternary
studies (abs.): Geological Society of America, Abstracts with
Programs, 20(7): A374.
Hawley, J. W. (ed.), 1978, Guidebook to Rio Grande rift in New Mexico
and Colorado: New Mexico Bureau of Mines & Mineral Resources,
Circular 163: 241 pp.
Ingersoll, R. V., Cavazza, W., Baldridge, W. S., Shafiqullah, M., 1990,
Cenozoic sedimentation and paleotectonics of north-central New Mexico: Implications for initiation and evolution of the Rio Grande rift:
Geological Society of America, Bulletin, 102: 1280-1296.
Kelley, V. C., 1979, Geomorphology of Espanola basin: New Mexico
Geological Society, Guidebook 7: 281-288.
Kelson, K. I., 1986a, Long-term tributary adjustments to base-level lowering northern Rio Grande rift, New Mexico: Unpublished MS thesis,
University of New Mexico, Albuquerque, 210 pp.
Kelson, K. I., 1986b, Long-term tributary adjustments to base-level lowering, northern New Mexico (abs.): Geological Society of America,
Abstracts with Programs, 18(5): 366.
Kelson, K. I., and Wells, S. G., 1987, Present-day fluvial hydrology andlong-term tributary adjustments, northern New Mexico; in Menges,
C. M. (ed.), Quaternary tectonics, landform evolution, soil chronologies
and glacial deposits—northern Rio Grande rift of New Mexico: Friends of
the Pleistocene Field-trip Guidebook, pp. 95-109.
Lambert, P. W., 1968, Quaternary stratigraphy of the Albuquerque area,
New Mexico: Unpublished PhD dissertation, University of New
Mexico, Albuquerque, 329 pp.
Love, D. W., Reimer, R. F., Hawley, J. W., Johnpeer, G. D., and Bobrow,
D. J., 1987, Summary of geotechnical investigations near Espanola,
New Mexico: Friends of the Pleistocene–Rocky Mountain cell, Fieldtrip guidebook, pp. 133-157.
Manley, K., 1976a, The late Cenozoic history of the Espanola basin, New
Mexico: Unpublished PhD dissertation, University of Colorado, Boulder, 171 pp.
Manley, K., 1976b, K–Ar age determinations of Pliocene basalts from the
Espanola basin, New Mexico: Isochron/West, no. 16: 29-30.
Manley, K., 1978, Pliocene deformation along the Rio Grande rift in the
Espanola basin, New Mexico (abs.): Geological Society of America,
Abstracts with Programs, 5(10): 233.
Manley, K., 1979a, Stratigraphy and structure of the Espanola basin, Rio
Grande rift: Tectonics and magmatism: American Geophysical Union,
Washington, D.C., pp. 71-86.
Manley, K., 1979b, Tertiary and Quaternary stratigraphy of the northeast
plateau, Espanola basin, New Mexico: New Mexico GeologicalSociety, Guidebook 30: 231-236.
Manley, K., 1982, Geologic map of the Cannes quadrangle, Rio Arriba
County, New Mexico: U.S. Geological Survey, Miscellaneous Field
Studies Map MF- 1440, scale 1:24,000.
Manley, K., and May, S. J., 1984, Third-day road log from Taos to Tres
Piedras, Tusas, Spring Canyon, Tres Piedras, Las Tablas, Petaca, La
Madera, and Ojo Caliente: New Mexico Geological Society,
Guidebook 35: 349-376.
Manley, K., and Mehnert, H. H., 1981, New K–Ar ages for Miocene andPliocene volcanic rocks in the northwestern Espanola basin and theirrelationships to the history of the Rio Grande rift: Isochron/West, no.
30: 5-8.
Menges, C. M., 1988, The tectonic geomorphology of mountain front
landforms in the northern Rio Grande rift, New Mexico: Unpublished
PhD dissertation, University of New Mexico, Albuquerque.
Morgan, P., and Golombek, M. P., 1984, Factors controlling the phases
and styles of extension in the northern Rio Grande rift: New Mexico
Geological Society, Guidebook 35: 13-19.
Muehlberger, W. R., 1979, The Embudo fault between Pilar and Arroyo
Hondo, New Mexico: An active intracontinental transform fault: New
Mexico Geological Society, Guidebook 30: 77-82.
Ruhe, R. V., 1967, Geomorphic surfaces and surficial deposits insouthern New Mexico: New Mexico Bureau of Mines & Mineral
Resources, Memoir 18: 1-66.
Sama-Wojcicki, A. M., Morrison, S. D., Meyer, C. D., and Hillhouse, J.
W., 1987, Correlation of upper Cenozoic tephra layers between sediments of the western United States and comparison withbiostratigraphic and magnetostratigraphic data: Geological Society ofAmerica, Bulletin, 98: 207-223.
Schumm, S. A., 1985, Explanation and extrapolation in geomorphology;
seven reasons for geologic uncertainty: Japanese Geomorphological
Union, Transactions, 6(1): 1-18.
Smith, C. T., Budding, A. J., and Pitrat, C. W., 1961, Geology of thesoutheastern part of the Chama basin: New Mexico Bureau of Mines &Mineral Resources, Bulletin 75, 57 pp.

45
Smith, H. T. U., 1938, Tertiary geology of the Abiquiu quadrangle, New
Mexico: Journal of Geology, 46: 933-965.
Smith, R. L., Bailey, R. A., and Ross, C. S., 1970, Geologic map of theJemez Mountains, New Mexico: U.S. Geological Survey,
Miscellaneous Geological Investigations Map 1-571, scale 1:25,000.
Spell, T. L., Harrison, T. M., and Wolff, J. A., 1989, 40Ar/39Ar dating of
the Bandelier Tuffs and associated ignimbrites: Constraints on
evolution of the Bandelier magma system: EOS, American GeophysicalUnion, Transactions, 70(43): 1413.
Vernon, J. H., and Riecker, R. E., 1989, Significant Cenozoic faulting,east margin of the Española basin, Rio Grande rift, New Mexico: Geology, 17: 230-233.
Waresback, D. B., 1986, The Puye Formation, New Mexico: Analysis ofa continental rift-filling volcaniclastic alluvial-fan sequence: Unpublished MS thesis, University of Texas at Arlington, 269 pp.
Waresback, D. B., and Turbeville, B. N., 1990, Evolution of a PlioPleistocene volcanogenic-alluvial fan: The Puye Formation, Jemez
Mountains, New Mexico: Geological Society of America, Bulletin,
102: 298-314.
Wells, S. G., Kelson, K. I., and Menges, C. M., 1987, Quaternary evolutionof fluvial systems in the northern Rio Grande rift, New Mexico andColorado: Implications for entrenchment and integration of drainage
systems; in Menges, C. M. (ed.), Quaternary tectonics, landform evolution, soil chronologies and glacial deposits—northern Rio Grande

rift of New Mexico: Friends of the Pleistocene, Field-trip Guidebook,pp. 55-69.
Wells, S. G., McFadden, L. D., and Schultz, J. D., 1990, Eolian landscapeevolution and soil formation in the Chaco dune field, southern
Colorado Plateau, New Mexico: Geomorphology, 3: 517-546.
Wilson, J. E., 1977, Stratigraphy and structure of part of the Abiquiu
embayment of the Rio Grande rift, New Mexico: Unpublished MS
thesis, University of New Mexico, Albuquerque, 89 pp.
Wright, H. E., Jr., 1946, Tertiary and Quaternary geology of the lower
Rio Puerco area, New Mexico: Geological Society of America,
Bulletin, 57: 383-456.
Wood, G. H., Jr., and Northrop, S. A., 1946, Geology of the Nacimiento
Mountains, San Pedro Mountains, and adjacent plateaus in parts of
Sandoval and Rio Arriba Counties, New Mexico: U.S. Geological Survey, Oil and Gas Investigations Preliminary Map 57.
Zoback, M. L., and Thompson, G. A., 1978, Basin and Range rifting in
northern Nevada: Clues from a mid-Miocene rift and its subsequent
offsets: Geology, 6: 111-116.
Zoback, M. L., and Zoback, M., 1980, State of stress in the conterminous
United States: Journal of Geophysical Research, 85: 6113-6156.
Zoback, M. L., Anderson, R. E., and Thompson, G. A., 1981, Cainozoic
evolution of the state of stress and style of tectonics of the Basin andRange province of the western United States: Royal Society of London,
Philosophical Transactions (A), 300: 407-434.

Appendix: Field-trip stops and topics of discussion
A detailed road log is Open-file Report no. 374, available
from New Mexico Bureau of Mines & Mineral Resources,
Socorro, New Mexico 87801.
Stop 1. Espanola High School. (1) Age of deposits behindhigh school. (2) Methods of Quaternary dating used in theregion.
Stop 2. Arroyo de la Presa. (1) Importance of the Embudo
fault as an intra-basin structure. (2) Discuss the changes inbasin evolution across the Embudo: relative base-levelstability north of the transfer fault and substantial Pliocene
aggradation south of the fault (refer to Fig. 4B and 4C).
Stop 3. Arroyos between Medanales and Abiquiu. (1)Traverse a flight of axial-stream Quaternary terraces (refer toFig. 4E). (2) Visit a deposit of Lava Creek B ash. (3) Discuss
landform evolution in rift (i.e. the proclivity for landscapeinversion).
Stop 4. Abiquiu Dam. View Quaternary surfaces formedin the Colorado Plateau and contrast with the landform evolution in the rift. (2) Entertain a discussion on the problems/difficulties with using soils on older Quaternary surfaces.
Overnight at Ghost Ranch.
Stop 5. Roadcut on US-84 near Ghost Ranch turnoff.
(1) Snail locality and use of amino-acid racemization dating
method. (2) Overlook Colorado Plateau and margin of rift tosee offset of the T1 surface (Cerro Pedernal, Mesa Escoba,

Polvadera Mesa, and Cañons Mesa). (3) Discuss Eocene tolate Miocene rift history.
Stop 6. US-84, 5 mi west of Abiquiu: Cañones fault zone.
(1) View the Cannes fault zone, which offsets upper Paleozoic
and Mesozoic rocks (to the west) against Tertiary rift-fillsediment (to the east) at the surface. (2) Gain anotherperspective on the offset of the T1 surface. (3) Finally view
offset on the Garcia fault zone.
Stop 7 (optional). Polvadera Mesa and overlook ofCañoncito Seco. Long drive over dirt roads culminating
in a lunch spot up in the ponderosa pines. (1) Cañoncito
Seco illustrates the magnitude of the pre-Bandelier
incision episode and the Bandelier aggradation episode(refer to Fig. 4D) and the post-Bandelier incision episode(refer to Fig. 4E).
Stop 8. Madera Canon pediments. (1) View displacement of Quaternary surfaces by the Madera Cañon fault
zone (see Fig. 6). (2) Look at soil pits across the fault
zone if possible.
Stop 9. White Rock Canyon overlook. Discuss: (1) the
Puye aggradation episode (refer to Fig. 4C); (2) west-tilting ofEspañola basin against Pajarito fault; (3) the pre-Bandelier
incision episode and Bandelier aggradation episode preservedin the paleocanyon across the Rio Grande; (4) the age ofgravel deposit (Qoal) over the Cerros del Rio basalt; and (5)wrap up the field trip.

47

New Mexico Bureau of Mines & Mineral Resources Bulletin 137, April 1991

Correlation of Triassic strata of the Colorado Plateau andsouthern High Plains, New Mexico
Spencer G. Lucas
New Mexico Museum of Natural History, P.O. Box 7010, Albuquerque, New Mexico 87194

Introduction
Triassic strata on the Colorado Plateau pertain to the
Moenkopi and Chinle Formations of Early—Middle andLate Triassic age, respectively. On the southern High Plainsof eastern New Mexico and west Texas, Triassic strata havetraditionally been assigned to the Dockum Group of Late
Triassic age. However, recent studies have identified the
Moenkopi Formation (Anton Chico Member) as the oldest
Triassic strata in eastern New Mexico (Lucas and Morales,
1985; Lucas and Hunt, 1987, 1989b).
Moenkopi, Chinle, and Dockum strata are nonmarine
redbeds deposited in fluvial and lacustrine environments.
Their chronology largely rests on vertebrate fossils and palynology (Lucas and Hunt, 1989a). Correlation of these
Triassic strata from the Colorado Plateau to the southernHigh Plains has long been uncertain, partly because of aperceived absence of Triassic strata across the Rio Grande
rift of central New Mexico (e.g. McKee et al., 1959; Finch
and Wright, 1983). However, significant outcrops ofTriassic strata are present in the Rio Grande rift of northcentral New Mexico, especially in the Hagan basin between
Albuquerque and Santa Fe (Fig. 1).

Recent research on the stratigraphy and paleontology of
the Triassic rocks across northern New Mexico allows relatively certain correlation of Moenkopi, Chinle, andDockum rocks from the Colorado Plateau to the southernHigh Plains. This correlation suggests continuity of Triassic
deposition across a broad region of the Southwest andundermines the proposed existence of a Pedernal uplift
separating Chinle and Dockum depositional basins during
the Late Triassic. Here, I review Triassic stratigraphy,
paleontology, and correlation from the eastern edge of theColorado Plateau to the southern High Plains and discuss
briefly the revised Triassic paleogeography mandated bythis review. The last section of this article presents a field
guide to Triassic outcrops in central New Mexico critical to
the review presented here.
Colorado Plateau edge: Lucero uplift to Chama basin
Triassic stratigraphy and paleontology from the Lucero uplift
to the Chama basin, along the eastern edge of the Colorado
Plateau (Fig. 1), have been recently reviewed by Kurtz (1978),
Kurtz and Anderson (1980), Woodward (1987), Lucas andHayden (1989a, b), Hunt and Lucas (1989a, b,

48
1990a, b), Hunt et al. (1989a, b), and Dubiel (1989b).
The oldest Triassic strata along this transect (Figs. 1, 2)
pertain to the Anton Chico Member of the Moenkopi Formation, as much as 68 m of grayish-red and reddish-brownlithic wackestones, mudstones, and siltstones that yieldcharophytes, ostracodes, bones of capitosauroid amphibians,
and large footprints of Chirotherium reptiles indicative of an
Early or Middle Triassic age (Kietzke, 1989a, b; Lucas andHayden, 1989a, b). These strata, which rest disconformablyon the Permian Glorieta and San Andres Formations, were
earlier identified as Permian Bernal Formation in the Jemez
Mountains (e.g. Wood and Northrop, 1946; Baars, 1962;
Woodward, 1987).
Upper Triassic strata that rest disconformably on the
Moenkopi Formation along the eastern edge of the Colorado
Plateau pertain to the Chinle Formation. Two sets of memberlevel nomenclature are used for Chinle strata (Fig. 2). In the
Lucero uplift, the Chinle consists of five members recognizedacross much of the southern edge of the Colorado Plateau (inascending order): (1) Shinarump Member, as much as 24 m ofsilica-pebble conglomerate/sandstone and pedogenicallymodified siltstone, mudstone, and sandstone, the latter
referred to informally as the "mottled strata"; (2) Bluewater
Creek Member, 100 m or more of red sandstone andmudstone; (3) more than 1000 m of Petrified Forest Member,
divided into lower and upper mudrock-dominated intervals by
the Sonsela Sandstone Bed and containing the Correo
Sandstone Bed at or near its top; (4) Owl Rock Member,about 36 m of grayish-red and very pale-green mudstone,grayish-red—purple bioturbated sandy siltstone and silty
sandstone; and (5) Rock Point Member, about 70 m oflaterally extensive beds of massive, moderately reddish-brownsandy siltstone and laminar to trough-crossbedded,
moderately reddish-brown, very fine-grained, mature,quartzose sandstone. Contrary to some previous work
(O'Sullivan, 1974, 1977; Dubiel, 1989), which only identified the Owl Rock and Rock Point Members as far east as
Thoreau and Lupton, respectively, these units are present as
far east as Petaca Pinta (T6N, R6W) in southeastern Cibola
County (Maxwell, 1988a, b; Lucas and Hunt, 1991). The
Middle Jurassic Entrada Sandstone overlies the Chinle
Formation in the Lucero uplift.
To the north, in the Jemez and Nacimiento Mountains, theChinle member-level terminology is different (Fig. 2), in partdue to lithofacies changes and in part due to a local
nomenclature developed by Huene (1911) and Wood andNorthrop (1946). The oldest Chinle strata are the mottled
strata and Agua Zarca Member (includes the "sandstone
member" of Stewart et al., 1972; cf. Woodward, 1987), asmuch as 35 m of tan and gray quartzose sandstone and silicapebble conglomerate. The Agua Zarca interfingers with theoverlying Salitral Member, mostly blue, green, and graybentonitic mudstone and siltstone as much as 30 m thick. The
overlying Poleo Member is as much as 49 m of strata
dominated by medium-grained sandstone and limestonepebble conglomerate which rests unconformably on the Salitral. As much as 224 m of dominantly red and purple
mudstone, sandstone, conglomerate, and pedogenic calcrete,
the petrified Forest Member, overlie the Poleo. Locally, theRock Point Member (siltstone member of Stewart et al.,1972, and Dubiel, 1989b) disconformably overlies the
Petrified Forest Member. It consists of as much as 70 m oflaterally extensive beds of moderately reddish-brownsiltstone and sandstone. The Middle Jurassic Entrada Sandstone overlies Triassic strata in the Jemez Mountains—
Chama basin area.
Correlation of the Triassic section in the Chama basinJemez Mountains with the Triassic section in the Lucero
uplift is readily made on the basis of lithology, paleontology,
and stratigraphic sequence. Thus, the Anton Chico Memberof the Moenkopi Formation is present in both areas, andcorrelation of the Shinarump and the Agua Zarca has long

49
been accepted (e.g. Wengerd, 1950). The small vertebrate
fauna of the Salitral Member includes large metoposaurs,
phytosaurs, the aetosaur Longosuchus, and a ?dinosaur(Hunt and Lucas, 1989b, 1990b). This late Carnian fauna isapproximately of the same age as the fauna of the
Bluewater Creek Member in west-central New Mexico,
which includes the phytosaur Angistorhinus. Both faunas
are of Paleorhinus-biochron age or early–late Carnian(Hunt and Lucas, in press).
In the Lucero uplift, the lower part of the Petrified Forest
Member has produced large and small metoposaurs, phytosaurs, and the aetosaur Stagonolepis (Hunt et al. , 1989a).
This fauna indicates a late Carnian age slightly younger than
the Paleorhinus biochron (Lucas, 1990; Hunt and Lucas,
1990b, in press). The Petrified Forest Member has two distinct faunas in the Jemez Mountains: an older fauna correlative with that of the lower Petrified Forest Member in the
Lucero uplift and a younger fauna equivalent to that of theupper Petrified Forest Member on the southern Colorado
Plateau (Hunt and Lucas, 1990a). The Rock Point Member in
the Chama basin contains the mass-death assemblage of theLate Triassic dinosaur "Coelophysis" at Ghost Ranch(Colbert, 1989; Hunt and Lucas, 1989a, b), but at present nobiochronologically useful fossils are known from the Rock
Point elsewhere on the Colorado Plateau.
Rio Grande rift: Hagan basin
The Triassic section exposed in the Hagan basin (Fig. 1)
is, with minor differences, very similar to that exposed in
the Jemez Mountains (Fig. 2). An essentially identical
Triassic section, much complicated by local structure, is
exposed in the Placitas area (Fig. 1), 5-10 km southwest of
the Hagan basin Triassic outcrops (Menne, 1989a, b).
The gap in Triassic exposure between the Jemez Mountains and Hagan basin is approximately 45 km. However,logs of two wildcat oil wells in this gap Shell Oil Company
Santa Fe Pacific No. 1 (Figs. 1, 3) and Avila Oil Company
Odium Federal No. 1—provide stratigraphic data that indicate continuity of the Triassic section from the Jemez
Mountains to the Hagan basin (Black and Hiss, 1974;
though my picks of the unit contacts on these logs differ
slightly from theirs). Thus, the Triassic section in the Shell
well begins with 4.3 m (14 ft: 8870-8884 ft) of Anton Chico
Member of the Moenkopi Formation resting on the Permian
Glorieta Sandstone. The overlying 43 m (140 ft: 8730-8870
ft) of sandstone of the Chinle Agua Zarca Member are followed by 348 m (1142 ft: 7728-8870 ft) of mudstone andvery minor sandstone of the Petrified Forest Member of the
Chinle Formation overlain by the Entrada Sandstone. The
uppermost Petrified Forest Member contains a 4.3 m (14 ft:
7736-7750 ft) thick sandstone that I identify as the Correo
Sandstone Bed. A similar sandstone interval occupies thesame stratigraphic position near the top of the exposed
Chinle Formation along the escarpment southwest of NM44 on the Zia Indian Reservation (T16N, R1W/R1E) in
Sandoval County.
The Avila well reveals a similar stratigraphic section: 36 m
(120 ft: 2495-2615 ft) of Moenkopi, 59 m (195 ft: 23002495
ft) of Agua Zarca, and 411 m (1350 ft: 950-2300 ft) of
Petrified Forest Member. However, the Correo Sandstone
Bed cannot be identified in the log, so it apparently is absent.
Furthermore, the difference in the thicknesses (4.3 m vs. 36m) of the Moenkopi Formation penetrated by the two wells
reflects the unconformity at the base of the overlying

Chinle Formation. Lucas and Hayden (1989a) demonstrated
a similar unconformity in outcrop between Red Mesa andGuadalupita Mesa in the Jemez Mountains.
Previous workers (Harrison, 1949; Reynolds, 1954; Kelley
and Northrop, 1975; Picha, 1982; Woodward and Picha,
1989) divided the Triassic strata in the Hagan basin into
basal Santa Rosa Sandstone and overlying Chinle Formation, except Smith (1961) who used Dockum Group divided
into Santa Rosa Sandstone and "upper member." Strata of
the Permian Bernal Formation were identified between the
Santa Rosa and the Permian San Andres Formation.
The Triassic section in the Hagan basin begins with theAnton Chico Member of the Moenkopi Formation, as muchas 20 m of grayish-red lithic wackestones and siltstones that
rest disconformably on the San Andres Formation (Fig. 4A;
Lucas and Hayden, in press). Kelley and Northrop (1975, p.
52) first identified these strata as Bernal Formation andstated that "locally the limestone of the Bernal contains
small brachiopods." I have not been able to verify this claim
and suspect that the brachiopods are from San Andres Limestone and were located as float on Moenkopi strata in this
topographically rugged and structurally complex area.
The overlying lowermost part of the Chinle Formation is as
much as 90 m of mostly laminar, grayish-white, quartzose
sandstone and silica-pebble conglomerate with fossil logs
(Fig. 4B, C) that have long been termed Santa Rosa Sandstone. However, I assign these strata to the Agua Zarca
Member of the Chinle Formation because they closely resemble the Agua Zarca Member in lithology, occupy an
identical stratigraphic position, and lack the medial mudstone interval (Los Esteros Member) characteristic of the
Santa Rosa Formation near Lamy, the closest outcrops assigned with certainty to the Santa Rosa. The well-log datasummarized above also support continuity of the Agua Zarca
Member of the Chinle Formation between the Jemez Mountains and the Hagan basin.
Indeed, in the Hagan basin the mudstone-dominated strata
immediately above the sandstones identified as Agua Zarca
are mostly blue, green, and gray bentonitic mudstone similar
to the Salitral Member to the west. Large metoposaur fossils
from limestone-pebble conglomerates in this interval (Fig.
4D) also lend some support to this correlation. However,since the Poleo Sandstone Member is absent in the Hagan
basin, I do not apply the name Salitral here. Instead, I refer
the entire mudstone-dominated section above the Agua Zarca
Member in the Hagan basin, the remainder of the Chinle
Formation, to the Petrified Forest Member.
This unit is as much as 410 m of reddish-brown bentonitic
mudstone intercalated with thin beds of laminar sandstone
and limestone-cobble conglomerate (Fig. 4E). The upper
part of the Petrified Forest Member is a prominent, cliffforming bed, 5.5-24 m thick, of crossbedded quartzose
sandstone and lithic- and limestone-pebble conglomerate(Figs. 4E–F, 5). These strata contain fragmentary remains ofphytosaurs and metoposaurs indicative of a Late Triassicage (Fig. 6). Their stratigraphic position, lithology, andpaleontology resemble closely the Correo Sandstone Bed of
the Petrified Forest Member of the Chinle Formation at its
type locality on Mesa Gigante in the Lucero uplift, 80 km to
the southwest (Lucas et al., 1987). For this reason, I assign
the Hagan basin strata to the Correo (Lucas et al., 1988). In
the Hagan basin, the medial siltstone member of the Middle
Jurassic Entrada Sandstone rests directly on the Correo(Figs. 4F, 5) or on a thin interval of upper Petrified

Forest Member mudstone above the Correo (Lucas et al.,
1988; Pigman and Lucas, 1989).
Just east of the Hagan basin, small, structurally complex
outcrops in the San Pedro Mountains—Cerrillos area ofwestern Santa Fe County (e.g. Read and Andrews, 1944;
Disbrow and Stoll, 1957; Bachman, 1976), poorly exposed
Triassic strata in the Cedar Crest—Sandia Park area ofnortheastern Bernalillo County (Kelley and Northrop, 1975),
and the log of a wildcat oil well north of Cerrillos (sec. 22,
T15N, R8E) provide the only data on the Triassic strata westof a 20-25 km wide gap of Triassic outcrop. Published work
just cited and unpublished data I have gathered indicate thatthese Triassic strata are essentially the same as those
exposed in the Hagan basin to the west, with one significant
difference: the Correo Sandstone Bed is not present.
Absence of the Correo is readily verified in roadcuts along
NM-14 north of Cedar Crest (sec. 2, T10N, R5E) and where
Arroyo del Tuerto crosses the highway just north of Golden.
At both locales, the Entrada Sandstone rests on a thick,reddish-brown mudstone section of the Petrified Forest
Member of the Chinle Formation. Indeed, the apparent
eastward pinch-out of the Correo Sandstone Bed occurs inthe Hagan basin just north of Puertecito (SW1/4 sec. 10,
T12N, R6E) where sandstone and conglomerate of theCorreo thin dramatically and pinch out into mudstone of theupper part of the Petrified Forest Member (Fig. 4F).
Eastern edge of Basin and Range: Lamy area
Triassic strata in the Lamy area of Santa Fe County havelong been assigned to the Dockum Group (Read et al., 1944;
Stearns, 1953; Dane and Bachman, 1965). Johnson (1973),
however, assigned these rocks to the Santa Rosa and Chinle
Formations, producing a stratigraphic nomenclature identical to that long used in the Hagan basin.
Recent work by B. Allen, P. Birchoff, A. Hunt, and myself,
however, indicates that the Triassic section exposed in the
Lamy area is very similar to that exposed in the Santa
Rosa—Tucumcari region of east-central New Mexico (Allenand Lucas, 1988). Thus, the base of the Triassic section near
Lamy is the Anton Chico Member of the Moenkopi

Formation, as much as 37 m thick and composed of grayishred, trough-crossbedded litharenite and conglomerate. It
contains bones of capitosauroid amphibians indicative of a
Middle Triassic age. South of Lamy, the Moenkopi rests
disconformably on moderately reddish-brown sandy siltstone of the Permian Bernal Formation (T 1 2-14N, R1011E). In northern Torrance County, west of Clines Corners
along 1-40 (secs. 13, 14, T9N, R11E), the Anton Chico
Member of the Moenkopi also is present between the Bernaland Santa Rosa Formations.
The Santa Rosa Formation rests disconformably on theAnton Chico Member in the Lamy area. It is as much as
140 m thick, and the three members recognized by Lucasand Hunt (1987; also see Gorman and Robeck, 1946) in
Guadalupe County can also be identified in the Lamy area:
basal, sandstone-dominated Tecolotito Member (as muchas 61 m thick), medial, mudstone-dominated Los Esteros
Member (as much as 47 m thick), and upper, sandstonedominated Tres Lagunas Member (as much as 62 m thick)(Hock, 1970; Johnson, 1973; Allen and Lucas, 1988; Huntand Lucas, 1988).
Overlying mudstones pertain to the Garita Creek Formation. The variegated blue, gray, green, and red bentonitic
mudstones and minor sandstones and limestone-pebble conglomerates are at least 40 m thick. The mass-death assemblage of Metoposaurus fraasi (Lamy amphibian quarry)from the Lamy area (Romer, 1939; Colbert and Imbrie,
1956; Hunt and Lucas, 1989b) is from the Garita Creek
Formation south of Lamy (SW1/4 SW1/4 NE 1/4 NE 1/4
sec. 29, T12N, R11E).
The Trujillo Formation rests disconformably on the Garita
Creek and consists of laminar and crossbedded sandstone andlimestone-pebble conglomerate 30-40 m thick. It is well
exposed along NM-41 south of the amphibian quarry where
it caps escarpments. To the west, near White Bluffs, about 40
m of the Bull Canyon Formation, dominantly reddish-brown
bentonitic mudstone, overlie the Trujillo. The Triassic section
south of Lamy is culminated by 20-25 m of laterally
persistent beds of moderately reddish-brown, fine-grainedsandstone, siltstone, and non-bentonitic mud-stone under the
Entrada Sandstone at White Bluffs. The stratigraphic positionand lithology of these strata are identical to the Redonda
Formation of east-central New Mexico, to which they are
assigned.
Southern High Plains: Santa Rosa to Tucumcari
Triassic stratigraphy and paleontology in the Santa RosaTucumcari region of east-central New Mexico have beenrevised recently by Lucas et al. (1985), Lucas and Hayden(1987, 1989b, c), Lucas and Morales (1985), Finch (1988),
Ash (1988), and Hunt et al. (1988). The Triassic section
consists of the same stratigraphic units present in the Lamyarea, but here they are generally thicker.
Paleogeographic significanceThe stratigraphic correlations proposed here necessitate achange in current concepts of the paleogeography of central
New Mexico during the Triassic. These concepts (Fig. 7)
identify a positive area in central New Mexico, the Pedernal
uplift, that separated Triassic deposition in the Chinle basin
to the west from deposition in the Dockum basin to the east
(McKee et al., 1959; McGowen et al., 1983; Dubiel, 1989a).
Indeed, Chinle paleocurrents on the Colorado Plateau are
directed to the north, northwest, and west (e.g.

Stewart et al., 1972; Dubiel, 1989a), consistent with thepresence of an uplift to the east.
However, paleocurrents measured in Triassic strata on theHigh Plains of eastern New Mexico and west Texas are also
generally directed to the north, northwest, and west (e.g.
Caveau, 1960; Kiatta, 1960; Lupe, 1988; May, 1988; Hayden
and Lucas, 1988; DeLuca and Eriksson, 1989; May andLehman, 1989). Only a small number of paleocurrents in the
Santa Rosa Formation indicates flow to the east (McGowen etal., 1979; Granata, 1981). Yet, these few paleocurrents havebeen combined with the generally northwest—southeast

orientation of Triassic sand bodies in the subsurface to reconstruct a regional Late Triassic paleoslope down to the east
(McGowen et al., 1983, figs. 25-26). However, theorientations of the sand bodies are bidirectional indicators offlow and more likely suggest paleoflow to the northwest.
The close similarity of the Triassic section south of Lamyand in the Santa Rosa—Tucumcari area also argues against a
Pedernal uplift at the location of the late Paleozoic uplift inwhat is now northeastern Torrance County and southwestern
San Miguel County. Indeed, Triassic strata of the Moenkopi
and Santa Rosa Formations still preserved at thisbasin. Continue 11.4 mi on this road to a locked gate on aroad that enters from the right. This land is property of the
Diamond Tail Ranch, and permission to enter and a key tothe gate must be obtained from the ranch headquarters.
Unlock gate and proceed about 2.5 mi to a left fork in theroad (after the water gap, before the ghost town of Tejon).
Turn left and continue 0.5 mi. Stop and climb the ridge tothe southwest. Here, on the east flank of Montezuma
Ridge, the Permo-Triassic boundary in the Hagan basin iswell exposed. San Andres, Moenkopi, and Agua Zarca
strata (Fig. 4A–C) can be examined here.
Stop 2: Top of Triassic section, Hagan basin
Return to main unpaved road through Hagan basin.
Continue east on this road for 4.3 mi. Stop at the virtual
ghost town of Puertecito. The bluff to the east of Puertecito
well exposes the upper part of the Triassic section in the
Hagan basin—Petrified Forest Member of Chinle
Formation—overlain by the Jurassic Entrada Sandstone(Fig. 4F). Also note the eastward pinchout of the Correo
Sandstone Bed displayed here (cf. Fig. 5).

location do not show thinning, paleocurrents, or facies
changes that would suggest that an uplift was located there
during the Triassic (Hock, 1970).
Some stratigraphic changes do occur in the 20-25 km
gap in Triassic exposure between the outcrops south of
Lamy and those just west of the Hagan basin (Fig. 1).
These changes are:
1. The Los Esteros and Tres Lagunas Members of the
Santa Rosa Formation, present south of Lamy, disappear to
the west where the Agua Zarca Member probably is equivalent to the Tecolotito Member of the Santa Rosa Formation.
A similar pinchout of the Los Esteros and Tres Lagunas
Members also occurs in the Sangre de Cristo Mountains ofnortheastern New Mexico (Lucas et al., 1990).
2. The Trujillo Formation, present south of Lamy, isabsent to the west. I suggest it pinches out westward into
mudstones of the Petrified Forest Member of the Chinle
Formation. This indicates the Petrified Forest Member ismostly equivalent to the Garita Creek and Bull Canyon
Formations, a conclusion well supported paleontologically(Lucas and Hunt, 1989b).
3. The Redondo Formation, present south of Lamy, isabsent to the west. Erosion at the base of the Entrada Sandstone almost certainly accounts for this. Thus, the Entrada
rests on the much older Correo Sandstone Bed in the Hagan
basin, but locally rests on the Redonda-correlative Rock
Point Member of the Chinle Formation in the Lucero uplift.
These changes in the Triassic stratigraphic section between Lamy and the Hagan basin are few and do not suggest
separate Triassic deposition in the two areas. Both Moenkopi
and Chinle depositional systems thus were continuous fromthe Colorado Plateau through the southern High Plains.
Field-trip stops
Stop 1: Base of Triassic section, Hagan basin
Beginning at the intersection of 1-25 and 1-40 in Albuquerque, proceed north on 1-25 towards Santa Fe. Leave I25 at Exit 257 (San Felipe Pueblo) and turn right at the stop
sign. The unpaved road heading east leads into the Hagan

Stop 3: Triassic strata, Lamy area
Continue east 5.7 mi on unpaved road through Hagan
basin until junction with paved NM-344. Turn left on NM344 and proceed north a total of 20.1 mi through Golden andMadrid and past Cerrillos to intersection (on right) withunpaved Santa Fe County Road 42 which leads to Galisteo.
Turn right and follow the county road 9.5 mi to Galistec(intersection with paved NM-41). Turn left on NM-41 andcontinue 5.9 mi to junction with NM-285 just south of Lamy.
Turn right and proceed south on NM-285 for 14.7 mi to pulloff at crest of hill. Here, extensive, structurally complex
outcrops on the Bernal, Moenkopi, Santa Rosa, Garita Creek,
and Trujillo Formations can be examined in the vicinity of
the famous Lamy amphibian quarry.
Stop 4: Base of Triassic section, Clines Corners
Continue south on NM-285 for 20.0 mi to Clines Corners.
Turn right and proceed west 3.4 mi on 1-40. Stop on right ofinterstate to examine San Andres, Bernal, Moenkopi, andSanta Rosa Formations well exposed here on the axis of the
Paleozoic Pedernal uplift. After stop, continue west on I40
to Albuquerque (47.5 mi).
Acknowledgments
I thank B. Allen, P. Birchoff, S. Hayden, A. Hunt, K.
Martini, P. Sealey, and M. Space for assistance in the field
and the laboratory. A. Hunt, B. Kues, and J. Lorenz provided helpful reviews of the manuscript.
References
Allen, B. D., and Lucas, S. G., 1988, Triassic stratigraphy near Lamy,Santa Fe County, New Mexico: New Mexico Geology, v. 10, p. 64. Ash,
S. R., 1988, Fossil plants from the mudstone member of the Santa Rosa
Formation at the principal reference section: U.S. Geological Survey,
Bulletin 1804, pp. 19-27.
Baars, D. L., 1962, Permian System of Colorado Plateau: American Association of Petroleum Geologists, Bulletin, v. 46, pp. 149-218. Bachman,G. 0., 1976, Geologic map of the Madrid quadrangle, Santa Fe andSandoval Counties, New Mexico: U.S. Geological Survey, Map GQ1268, scale 1:62,500.
Black, B. A., and Hiss, W. L., 1974, Structure and stratigraphy in thevicinity of the Shell Oil Company Santa Fe Pacific No. 1 test well,southern Sandoval County, New Mexico: New Mexico Geological Society, Guidebook 25, pp. 365-370.

55
Cazeau, C. J., 1960, Cross-bedding directions in Upper Triassic sandstonesof West Texas: Journal of Sedimentary Petrology, v. 30, pp. 459-465.
Colbert, E. H., 1989, The Triassic dinosaur Coelophysis: Museum of
Northern Arizona, Bulletin 57, 160 pp.
Colbert, E. H., and Imbrie, J., 1956, Triassic metoposaurid amphibians:
American Museum of Natural History, Bulletin, v. 110, pp. 399-452.
Dane, C. H., and Bachman, G. 0., 1965, Geologic map of New Mexico:
U.S. Geological Survey, Denver, scale 1:500,000.
DeLuca, J. L., and Eriksson, K. A., 1989, Controls on synchronous
ephemeral- and perennial-river sedimentation in the middle sandstone
member of the Triassic Chinle Formation, northeastern New Mexico,
U.S.A.: Sedimentary Geology, v. 61, pp. 155-175.
Disbrow, A. E., and Stoll, W. C., 1957, Geology of the Cerrillos area,
Santa Fe County, New Mexico: New Mexico Bureau of Mines &Mineral Resources, Bulletin 48, 73 pp.
Dubiel, R. F., 1989a, Depositional and climatic setting of the Upper Triassic Chinle Formation, Colorado Plateau; in Lucas, S. G., and Hunt, A. P.
(eds.), Dawn of the age of dinosaurs in the American Southwest: New
Mexico Museum of Natural History, Albuquerque, pp. 171-187.
Dubiel, R. F., 1989b, Depositional environments of the Upper Triassic
Chinle Formation in the eastern San Juan Basin and vicinity, New Mexico: U.S. Geological Survey, Bulletin 1808, pp. 1-22.
Finch, W. I., 1988, Description of the principal reference section for
the Santa Rosa Formation: U.S. Geological Survey, Bulletin 1804,
pp. 110.
Finch, W. I., and Wright, J. C., 1983, Measured stratigraphic sections ofuranium-bearing Upper Triassic rocks of the Dockum basin, eastern
New Mexico, West Texas, and the Oklahoma Panhandle with brief
discussion of stratigraphic problems: U.S. Geological Survey, Open-file
Report 83701, 118 pp.
Gorman, J. M., and Robeck, R. C., 1946, Geology and asphalt deposits ofnorth-central Guadalupe County, New Mexico: U.S. Geological Survey,
Oil and Gas Investigations Preliminary Map 44.
Granata, G. E., 1981, Regional sedimentation of the Late Triassic
Dockum Group, West Texas and eastern New Mexico: UnpublishedMA thesis, University of Texas at Austin, 199 pp.
Harrison, E. P., 1949, Geology of the Hagan coal basin: Unpublished MS
thesis, University of New Mexico, Albuquerque, 82 pp.
Hayden, S. N., and Lucas, S. G., 1988, Stratigraphy of the Permo-Triassic
boundary in northern New Mexico: Abstracts of the 41st Symposium on
Southwest Geology and Paleontology [Museum of Northern Arizona,
Flagstaff], p. 5.
Hock, P. F., Jr., 1970, Effect of the Pedernal axis on Permian and Triassic
sedimentation: Unpublished MS thesis, University of New Mexico, Albuquerque, 51 pp.
Huene, F., von, 1911, Kurze Mitteilung Ober Perm, Trias und Jura in New
Mexico: Neues Jahrbuch fur Mineralogie, Geologie und Palãontologie,
Beilage-Band 32, pp. 730-739.
Hunt, A. P., and Lucas, S. G., 1988, Late Triassic fauna from the Los
Esteros Member of the Santa Rosa Formation, Santa Fe County, New
Mexico and its biochronological implications: New Mexico Journal of
Science, v. 28, pp. 107-116.
Hunt, A. P., and Lucas, S. G., 1989a, Stratigraphy and vertebrate biochronology of Upper Triassic strata in the Chama basin, north-central New
Mexico: Abstracts of the Symposium on Southwestern Geology andPaleontology [Museum of Northern Arizona, Flagstaff] 1989, p. 15.
Hunt, A. P., and Lucas, S. G., 1989b, Late Triassic vertebrate localities in
New Mexico; in Lucas, S. G., and Hunt, A. P. (eds.), Dawn of the ageof dinosaurs in the American Southwest: New Mexico Museum of
Natural History, Albuquerque, pp. 72-101.
Hunt, A. P., and Lucas, S. G., 1990a, Paleontology and biochronology of
the Petrified Forest Member of the Upper Triassic Chink Formationnear San Ysidro, Sandoval County, New Mexico: New Mexico Journalof Science, v. 30, pp. 17-26.
Hunt, A. P., and Lucas, S. G., 1990b, Re-evaluation of "Typothorax" meadei,a Late Triassic aetosaur from the United States: Palãontologische
Zeitschrift, v. 64, pp. 317-328.
Hunt, A. P., and Lucas, S. G. (in press), The Paleorhinus biochron andthe correlation of the nonmarine Upper Triassic of Pangaea: Palaeontology.
Hunt, A. P., Lucas, S. G., Martini, K., and Martini, T., 1989[a], Triassic
stratigraphy and paleontology, Mesa del Oro, Valencia County, New
Mexico: New Mexico Geological Society, Guidebook 40, pp. 8-9.
Hunt, A. P., Lucas, S. G., and Sealey, P. L., 1989[b], Paleontology andvertebrate biochronology of the Upper Triassic Garita Creek Formation,
east-central New Mexico: New Mexico Journal of Science, v. 29, pp.
61-68.
Johnson, R. B., 1973, Geologic map of the Bull Canyon quadrangle, SantaFe County, New Mexico: U.S. Geological Survey, Map GQ-l077, scale
1:24,000.
Kelley, V. C., and Northrop, S. A., 1975, Geology of the Sandia
Mountains and vicinity, New Mexico: New Mexico Bureau of Mines &Mineral Resources, Memoir 29, 135 pp.
Kiatta, H. W., 1960, A provenance study of the Triassic deposits of northwestern Texas: Unpublished MS thesis, Texas Tech University,
Lubbock, 63 pp.
Kietzke, K. K., 1989a, Calcareous microfossils from the Triassic of thesouthwestern United States; in Lucas, S. G., and Hunt, A. P. (eds.),
Dawn of the age of dinosaurs in the American Southwest: New Mexico
Museum of Natural History, Albuquerque, pp. 223-232.
Kietzke, K. K., 1989b, Calcareous microfossils from the Moenkopi Formation (Triassic, Scythian or Anisian) of central New Mexico: New
Mexico Geological Society, Guidebook 40, pp. 181-190.
Kurtz, D. D., 1978, Sedimentology and stratigraphy of the Triassic Chinle
Formation, eastern San Juan Basin, New Mexico: Unpublished MS
thesis, Rice University, Houston, 177 pp.
Kurtz, D. D., and Anderson, J. B., 1980, Depositional environments andpaleocurrents of Chinle Formation (Triassic), eastern San Juan Basin,
New Mexico: New Mexico Geology, v. 2, pp. 22-27.
Lucas, S. G., 1990, The rise of the dinosaur dynasty: New Scientist, no.
1737, pp. 44-46.
Lucas, S. G., and Hayden, S. N., 1989a, Triassic stratigraphy of westcentral New Mexico: New Mexico Geological Society, Guidebook 40,
pp. 191-211.
Lucas, S. G., and Hayden, S. N., 1989b, Middle Triassic Moenkopi
Formation, Nacimiento Mountains, north-central New Mexico; in Lorenz, J. C., and Lucas, S. G. (eds.), Energy frontiers in the Rockies:
Albuquerque Geological Society, Albuquerque, pp. 16-17.
Lucas, S. G., and Hayden, S. N. (in press), Type section of the Permian
Bernal Formation and the Permo-Triassic boundary in north-central
New Mexico: New Mexico Geology.
Lucas, S. G., and Hunt, A. P., 1987, Stratigraphy of the Anton Chico andSanta Rosa Formations, Triassic of east-central New Mexico: Journal of
the Arizona-Nevada Academy of Science, v. 22, pp. 21-33.
Lucas, S. G., and Hunt, A. P. (eds.), 1989a, Dawn of the age of dinosaurs
in the American Southwest: New Mexico Museum of Natural History,
Albuquerque, 414 pp.
Lucas, S. G., and Hunt, A. P., 1989b, Revised Triassic stratigraphy in
the Tucumcari basin, east-central New Mexico; in Lucas, S. G., andHunt, A. P. (eds.), Dawn of the age of dinosaurs in the American
Southwest: New Mexico Museum of Natural History, Albuquerque,
pp. 150-170.
Lucas, S. G., and Hunt, A. P. (in press), Upper Triassic Owl Rock andRock Point Members, Chinle Formation, Petaca Pinta, Cibola County,
New Mexico: New Mexico Geology.
Lucas, S. G., and Morales, M., 1985, Middle Triassic amphibian from thebasal Santa Rosa Formation, east-central New Mexico: New Mexico
Geological Society, Guidebook 36, pp. 56-58.
Lucas, S. G., Allen, B. D., and Hayden, S. N., 1987, Type section of theTriassic Correo Sandstone Bed, Chinle Formation, Cibola County, New
Mexico: New Mexico Journal of Science, v. 27, pp. 87-93.
Lucas, S. G., Hunt, A. P., and Huber, P., 1990, Triassic stratigraphy in the
Sangre de Cristo Mountains, New Mexico: New Mexico GeologicalSociety, Guidebook 41, pp. 305-318.
Lucas, S. G., Hunt, A. P., and Morales, M., 1985, Stratigraphic nomenclature and correlation of Triassic rocks of east-central New Mexico: apreliminary report: New Mexico Geological Society, Guidebook 36, pp.
171-184.
Lucas, S. G., Martini, K., and Martini, T., 1988, Upper Triassic Correo
Sandstone Bed, Petrified Forest Member, Chinle Formation, Hagan basin,
Sandoval County, New Mexico: New Mexico Geology, v. 10, p. 65.
Lupe, R., 1988, Interpretive depositional history of the Santa Rosa Formation at the principal reference section: U.S. Geological Survey, Bulletin 1804, pp. 11-18.
Maxwell, C. H., 1988a, Geologic map of the Cerro del Oro quadrangle,
Cibola County, New Mexico: U.S. Geological Survey, Miscellaneous
Field Studies Map MF-2033, scale 1:24,000.
Maxwell, C. H., 1988b, Geologic map of the Marmon Ranch quadrangle,
Cibola County, New Mexico: U.S. Geological Survey, Miscellaneous
Field Studies, Map MF-2049, scale 1:24,000.
May, B. A., 1988, Depositional environments, sedimentology, and stratigraphy of the Dockum Group (Triassic) in the Texas Panhandle: Unpublished MS thesis, Texas Tech University, Lubbock, 180 pp.
May, B. A., and Lehman, T. M., 1989, Sedimentology of the Dockum
Group (Triassic), Texas Panhandle (abs.): Geological Society of America, Abstracts with Programs, v. 21, p. 34.

56
McGowen, J. H., Granata, G. E., and Seni, S. J., 1979, Depositionalframework of the lower Dockum Group (Triassic) Texas Panhandle:
Bureau of Economic Geology, University of Texas, Report of Investigations 97-1979, 60 pp.
McGowen, J. H., Granata, G. E., and Seni, S. J., 1983, Depositionalsetting of the Triassic Dockum Group, Texas Panhandle and eastern
New Mexico; in Reynolds, M. W., and Dolly, E. D. (eds.), Mesozoic
paleogeography of the west-central United States: Rocky Mountain
Section SEPM, Denver, pp. 13-38.
McKee, E. D., Oriel, S. S., Ketner, K. B., MacLachlan, M. E. H.,
Goldsmith, J. W., MacLachlan, J. C., and Mudge, M. R., 1959, Paleotectonic maps of the Triassic System: U.S. Geological Survey, Miscellaneous Geological Investigations Map 1-300, 33 pp.
Menne, B., 1989a, Stratigraphy and structure of the northern end of theSandia uplift; in Lorenz, J. C., and Lucas, S. G. (eds.), Energy
frontiers in the Rockies: Albuquerque Geological Society,Albuquerque, pp. 23.
Menne, B., 1989b, Structure of the Placitas area, northern Sandia uplift,
Sandoval County, New Mexico: Unpublished MS thesis, University of
New Mexico, Albuquerque, 163 pp.
O'Sullivan, R. B., 1974, The Upper Triassic Chinle Formation in northcentral New Mexico: New Mexico Geological Society, Guidebook 25,
pp. 171-174.
O'Sullivan, R. B., 1977, Triassic rocks in the San Juan Basin of New
Mexico and adjacent areas: New Mexico Geological Society,
Guidebook 28, pp. 139-146.
Picha, M. G., 1982, Structure and stratigraphy of the Montezuma salient,
Hagan basin area, Sandoval County, New Mexico: Unpublished MS
thesis, University of New Mexico, Albuquerque, 248 pp.
Pigman, C., and Lucas, S. G., 1989, The Jurassic section in the Hagan
basin, Sandoval County, New Mexico; in Lorenz, J. C., and Lucas, S.
G. (eds.), Energy frontiers in the Rockies: Albuquerque GeologicalSociety, Albuquerque, pp. 3-5.
Read, C. B., and Andrews, D. A., 1944, Geology of a part of the upper
Pecos River and Rio Galisteo region, N. Mex.: U.S. Geological Survey,

Oil and Gas Investigations Preliminary Map 8.
Read, C. B., Wilpolt, R. H., Andrews, D. A., Summerson, C. H., andWood, G. H., 1944, Geologic map and stratigraphic sections ofPermian and Pennsylvanian rocks of parts of San Miguel, Santa Fe,
Sandoval, Bernalillo, Torrance and Valencia Counties, north-central
New Mexico: U.S. Geological Survey, Oil and Gas Investigations
Preliminary Map 21.
Reynolds, C. B., 1954, Geology of the Hagan-La Madera area, SandovalCounty, New Mexico: Unpublished MS thesis, University of New Mexico, Albuquerque, 82 pp.
Romer, A. S., 1939, An amphibian graveyard: Scientific Monthly, v. 49,
pp. 337-339.
Smith, C. T., 1961, Triassic and Jurassic rocks of the Albuquerque area:
New Mexico Geological Society, Guidebook 12, pp. 121-128.
Stearns, C. E., 1953, Tertiary geology of the Galisteo-Tonque area, New
Mexico: Geological Society of America, Bulletin, v. 64, pp. 459-508.
Stewart, J. H., Poole, F. G., and Wilson, R. F., 1972, Stratigraphy andorigin of the Chinle Formation and related Upper Triassic strata in the
Colorado Plateau region: U.S. Geological Survey, Professional Paper
690, 336 pp.
Wengerd, S. A., 1950, Triassic rocks of northwestern New Mexico andsouthwestern Colorado: New Mexico Geological Society, Guidebook 1,
pp. 67-75.
Wood, G. H., and Northrop, S. A., 1946, Geology of the Nacimiento
Mountains, San Pedro Mountain, and adjacent plateaus in parts of Sandoval and Rio Arriba Counties, New Mexico: U.S. Geological Survey,
Oil and Gas Investigations Map OM-57, scale 1:95,000.
Woodward, L. A., 1987, Geology and mineral resources of Sierra Nacimiento and vicinity, New Mexico: New Mexico Bureau of Mines &Mineral Resources, Memoir 42, 84 pp.
Woodward, L. A., and Picha, M. G., 1989, Tectonic setting and structureof the Hagan basin, north-central New Mexico; in Lorenz, J. C., andLucas, S. G., Energy frontiers in the Rockies: Albuquerque GeologicalSociety, Albuquerque, pp. 167-169.

New Mexico Bureau of Mines & Mineral Resources Bulletin 137, April 1991

57

Tectonics, intrusive rocks, and mineralization of the San Pedro—Ortiz
porphyry belt, north-central New Mexico
Stephen R. Maynard', Lee A. Woodward', and David L. Giles³'LAC Minerals, USA, Inc., P.O. Box 21390, Reno, Nevada 89515; 'Department of Geology, University of New Mexico, Albuquerque, New Mexico 87131;

Introduction
Acknowledgments
The purpose of this field trip is to examine two major
tectonic and igneous features of north-central New Mexico,
The permission of Pegasus Gold, LAC Minerals, USA,the Tijeras—Canoncito fault system (TCFS) and the San Inc., Placer Dome Mining, and Cerrillos Gravel Products to
Pedro—Ortiz porphyry belt and their relationship to Tertiary enter their properties and publish information regardingbase-and precious-metal mineralization.
geology and mineralization is gratefully acknowledged.
Day 1 will be spent viewing the southern end of the Many geologists, too numerous to name individually, have
Tijeras—Canoncito fault system, the San Pedro Mountains, contributed to our understanding of the geology of the San
and the Cerrillos Hills. The group will return to Albuquerque Pedro—Ortiz porphyry belt. Alice Mueller drafted the figfor the evening. On Day 2 the trip will visit the Ortiz Moun- ures. Felix Mutschler and Tommy Thompson reviewed thetains, stopping at gold deposits or occurrences localized along manuscript, made corrections, and offered useful suggesthe Tijeras—Canoncito fault system.
tions.

Tectonic setting of the San Pedro—Ortiz porphyry belt
Lee A. Woodward
The San Pedro—Ortiz porphyry belt trends north-northeasterly, lying east of the Rio Grande rift and obliquely
straddling the northeast-trending Tijeras—Canoncito fault
system (Fig. 1). This belt of mid-Tertiary intrusions is in thenorth-trending New Mexico alkalic gold belt which includes
the Nogal—Bonito, Schelerville (West Bonito), White Oaks,
Jicarilla, Tecolote, and Gallinas mining districts. The New
Mexico alkalic gold belt is part of a larger belt of midTertiary gold-bearing alkalic intrusive centers that extends
from Montana to west Texas. Major tectonic elements near
the San Pedro—Ortiz porphyry belt include the Hagan andEspañola basins of the Rio Grande rift, the Sandia—
Manzano uplift, the Estancia basin, and the Tijeras—
Canoncito fault system.
The late Cenozoic Rio Grande rift is composed of several
north-trending grabens (commonly called basins) arranged
en-echelon north-northeasterly for a distance of about 725
km in New Mexico and Colorado (Kelley, 1952, 1977,1978). The Española basin, north of the Ortiz porphyry belt,
is 65 to 80 km long, 30 to 65 km wide, and is tilted mainly to
the west where it is bounded by high-angle faults. Within
this basin are antithetic and synthetic faults, forming secondorder horsts and grabens. The southeast margin of the basin
is marked by the Tijeras—Canoncito fault system and the
San Pedro—Ortiz porphyry belt.
The Hagan basin is an east-tilted half-graben that forms
an embayment or bench (Kelley, 1977) on the northeast
corner of the larger Albuquerque basin (Fig. 1). On the east,the Hagan basin is bounded by the La Bajada normal fault
which is downthrown on the west side. The western boundary of the basin is marked by the San Francisco normal fault
which is also downthrown on the west side. A broad dip-slope on the east side of the Sandia—Manzano upliftmerges through moderately dipping beds with the southernand southwestern parts of the Hagan basin. Maximum
structural relief between the deepest part of the Hagan
basin and the top of the Sandia uplift is about 6000 m. A
structure section that was based on seismic data (Black and
Hiss, 1974) across the northern part of the basin shows
about 4000 m of strati-graphic separation across the San
Francisco fault and 3300 m across the La Bajada fault.
Displacement along the latter dies out southward. La
Bajada and San Francisco are major faults that controlledlate Cenozoic sedimentation in this part of the Rio Grande
rift. Structures within the Hagan basin include low- andhigh-angle normal faults, high-angle reverse faults, Hagan
syncline, Espinaso monocline, and a few minor folds.
Arcuate outcrop patterns strike northerly on the west sideof the basin and swing eastward on the south margin, thus
defining the gently northeast-plunging Hagan syncline. The
Espinaso monocline trends north-northwesterly, faces east,and can be traced for about 10 km along strike.
The Sandia and Manzano Mountains form a north-trending, east-tilted fault block that is approximately 120 kmlong and up to 15 km wide. Precambrian rocks are exposedon the western fault scarp and the dip-slope to the east isdeveloped mainly on Pennsylvanian strata dipping about
15° (Kelley and Northrop, 1975). Maximum structuralrelief between the Sandia uplift and the northern part of the
Estancia basin is around 3000 m. Phanerozoic rocks in the
Ortiz porphyry belt also generally dip 15-25° east.
The western margin of the Estancia basin is transitionalwith the eastern dip-slope of the Manzano uplift. On the north,this basin is bounded by the Tijeras—Canoncito fault

system. There appears to be no major crustal extension
involved in the evolution of this basin. Rather it seems to haveformed in two stages, by Laramide and older rise of thePedernal uplift on the east and late Cenozoic eastward tiltingof the Sandia—Manzano uplift.
The Tijeras—Canoncito fault system consists of high-angle
faults that split and join along strike and form sigmoidal
patterns in map view, with horsts and grabens occurring
locally between major faults (Fig. 1). On its southwestern end,
the Tijeras fault has Precambrian metamorphic rocks

juxtaposed with Pennsylvanian strata; Quaternary alluvium
overlaps the bedrock outcrops to the southwest in the Albuquerque basin of the Rio Grande rift. To the northeast, theTijeras fault juxtaposes Precambrian rocks of contrasting
lithologies (Fig. 2) (Connolly, 1982) for nearly 10 km, withgreenstone and associated rocks on the southeast (Bruns,
1959) and quartzofeldspathic gneiss on the northwest (Lodewick, 1960). In the Tijeras mining district, the greenstone
hosts ore deposits that are interpreted as Precambrian syngenetic mineralization remobilized during later episodes ofmetamorphism and deformation (Fulp et al., 1982).
The next segment of the fault extends from the easternbase of the Sandia uplift to the northeast end of the Tijeras
graben. Rocks ranging in age from Pennsylvanian to Cretaceous are offset, and vertical separation along this segment
ranges from 0 to at least 1000 m (Kelley and Northrop,1975), with the southeast side down. Kelley and Northrop(1975, pp. 87-88) noted 500-600 m of left-lateral separationof fold axes and strata. The Gutierrez fault connects with the
Tijeras fault and extends about 24 km to the northeast,bounding the southeastern margins of the Tijeras graben andMonte Largo horst. Limited exposures indicate that the fault
is nearly vertical and has a maximum vertical separation of
about 1100 m where it juxtaposes Paleozoic and Mesozoic
rocks. To the northeast the fault splays toward the east andapparently dies out beneath alluvium. The northwesternboundary of Monte Largo horst is the continuation of the
Tijeras fault and has Pennsylvanian strata against
Precambrian metamorphic rocks, with about 1000 m of vertical separation (Kelley and Northrop, 1975). The Tijeras
fault continues to the northeast and curves eastward, forminga sigmoidal bend north of the San Pedro Mountains. The San
Pedro Mountains have many small faults that are parallel to
the Tijeras system.
In the Ortiz Mountains a graben, confined to the northeasttrending part of a sigmoidal bend, is present where bounding
faults displace Tertiary sedimentary strata and Tertiary intrusive rocks. There is about 1300 m of structural relief
between the Ortiz graben and the adjacent area to the south

east (Maynard et al., 1990). Northeast of the Ortiz Mountains
the fault is covered by upper Tertiary—Quaternary gravel(Bachman, 1975) but reappears at the surface where lower to
middle Tertiary strata are offset by the fault.
Lisenbee et al. (1979) have shown that the complex geometry along the fault system is the result of at least four
episodes of deformation, ranging from Precambrian to lateCenozoic. Direction of slip along any given segment may
have varied from one deformational episode to another;also, different segments may have undergone different slip
during the same episode of deformation. Precambrian
movement may have been largely strike-slip (Lisenbee etal. , 1979), establishing a major zone of anisotropy thatprobably extends through the crust and has had recurrent
movement. Late Cenozoic displacement was complex andhad a large component of dip-slip movement andsubordinate strike-slip. The composite strike component of
movement along the system from the Ortiz Mountains tothe southwest appears to have been left-slip and may be
responsible for dilation where faults change strike from
northeast to east-northeast. This could be the origin of the
Ortiz and Galisteo grabens. Thus, these features appear to
be of transtensional origin.
Movement during the late Cenozoic on the Tijeras—Canoncito system is interpreted to be that of a continental
transform, with extensional terrane to the northwest(Española and Hagan basins) and nonextended crust to thesoutheast (Estancia basin and Manzano dip-slope). Northerly trending extensional faults related to rifting are present

60
north of the Tijeras—Canoncito system in the Hagan andEspanola basins. These extensional and transform faults may
have controlled the trend and location of the Ortiz porphyry
belt. In a more detailed sense, faults and related fractures
have exerted structural control for mineral deposits by providing shear zones in Precambrian basement rocks, open
fractures for fissure filling, open-space breccias, and, of mostimportance, a zone of crustal weakness for magma
emplacement and conduits for hydrothermal fluids. The
magmas in turn appear to be genetically related to many ofthe mineral deposits, having been the source of mineralizing
solutions or having provided heat for hydrothermal systems.
A full discussion of the stratigraphy of the San Pedro—
Ortiz porphyry belt is beyond the scope of this paper. Stratigraphic relations within the belt have been extrapolated principally from exposures in the Hagan Basin (Black, 1979;
Picha, 1982) and are summarized in Fig. 3.
In general, the oldest supracrustal rocks in the San Pedro—
Ortiz porphyry belt are exposed at the southern end, in the San
Pedro Mountains and at South Mountain. The youngest
supracrustal rocks crop out at the northern end, in the Cerrillos
Hills.
Certain ore deposits and mineral occurrences in the San
Pedro—Ortiz porphyry belt are found in specific stratigraphic
contexts. At the San Pedro mine, copper-skarn mineralization
is localized in favorable beds of the Pennsylvanian Madera
Formation. In the Ortiz Mountains, the Greenhorn Limestone
Member of the Upper Cretaceous Mancos Shale hosts
copper—gold skarn at Lukas Canyon. Sills of latite porphyry,
whose emplacement was localized by Mancos and Menefee
Shales, host gold mineralization at the Carache Canyon
breccia pipe. And brecciated sandstone (quartzite) of the
Eocene Galisteo Formation contains gold and tungsten
mineralization at Cunningham Hill.
Road log: Day 1, Stop 1
Drive east on 1-40 from Albuquerque. About 3 mi after
entering Tijeras Canyon, round Deadman's Curve. East of
Deadman's Curve 1-40 is in the strike valley of the Tijeras
fault.

Stop 1. About 200 m east of Deadman's Curve on 1-40 inTijeras Canyon. Be sure to park well away from the traffic
lanes and stay on the shoulder.
Across the highway note the exposure of the Tijeras fault
where Quaternary colluvium is in fault contact with Proterozoic Tijeras greenstone (Fig. 2). Quartzite within the
Cibola gneiss makes up the bold outcrops on top of the nearridge to the north. To the south, Pennsylvanian Madera
Formation limestone holds up the prominent ridge. The
lower slopes of the ridge are underlain by Proterozoic Tijeras
greenstone.

Tijeras Canyon, Sandia Mountains, and Tijeras graben
Lee A. Woodward
Tijeras Canyon is underlain by the Tijeras fault which
strikes northeast and separates two Proterozoic metamorphic
terranes of contrasting rock types and structural styles (Connolly, 1981, 1982). The Tijeras fault has undergone at least
four episodes of movement, during the Precambrian, Pennsylvanian, Tertiary, and Quaternary (Lisenbee et al., 1979).
Although the fault may have had different directions of slip
during the various episodes of movement, it seems likelythat the cumulative sense of motion in Tijeras Canyon consists of several kilometers of left-slip as judged from contrasts in lithology, structural geometry, and metamorphic
history (Connolly, 1982).
Northwest of the Tijeras fault is the Cibola gneiss, a unit
derived mainly from arkose, with subordinate sandstone andshale protoliths. The gneiss is intruded by the Precambrian
Sandia granite and is characterized by augen gneiss near thecontact. Southeast of the fault the Tijeras greenstone consists
mostly of metabasalt with subordinate metasedimentary rocks,
including chlorite and biotite phyllite as well as metaquarzite.
Minor felsic metavolcanic rocks are also present. Both the
Cibola gneiss and Tijeras greenstone were isoclinally folded
during amphibolite-facies metamorphism, but the greenstone
was tightly refolded about axes nearly parallel to the older foldaxes, whereas the gneiss was refolded about axes at a highangle to the earlier fold axes (Connolly, 1982).
Ore deposits hosted by the Tijeras greenstone are interpreted as syngenetic, exhalative gold—copper mineralization
remobilized during metamorphism and deformation (Fulp etal., 1982). At the Great Combination and Mary M mines,
gold-bearing veinlets are associated with pyrite and malachite, and at the York mine, banded, semi-massive pyritechalcopyrite is present. Production from this district is small;

61
Elston (1967) estimated $10,000 of production from 1828 to1959 for the Tijeras and several nearby districts.
North of Tijeras Canyon, the dip-slope formed on Pennsylvanian strata on the east side of the Sandia Mountains isbroken by a mosaic of high-angle faults mostly having small
displacements and trending westerly, northwesterly, northerly, and northeasterly (Kelley and Northrop, 1975). Highangle reverse faults in the southeastern part of the Sandia
Mountains were suggested to be of Laramide age (Late
Cretaceous—early Tertiary) by Kelley and Northrop (1975).
In the absence of stratigraphic evidence to pin down such anage, it seems more likely that these reverse faults formed in
response to a local compressional stress field in the hinge or
synclinal bend of the eastward-tilted Sandia fault block in
the late Cenozoic.
The Tijeras graben is the southwestern part of anortheast-trending block bounded on the northwest by the
Tijeras fault and on the southeast by the Gutierrez fault.
This block, about 25 km long and up to 5 km wide, is tilted
to the southwest, with the Tijeras graben grading
northeastward into the Monte Largo horst. The boundary
between the graben and the horst is placed at the change
from downthrown to upthrown stratigraphic separation
across the block. Precambrian rocks are exposed in thehorst, and Cretaceous Mesaverde strata form the surface inthe central part of the graben. These latter rocks are coalbearing and were mined underground around 1900. Threecoal seams ranging from 0.3 to 0.7 m in thickness arepresent (Kelley and Northrop, 1975).
Maximum structural relief from the Tijeras graben
across the bounding fault on the southeast is approximately
1500 m. Within the graben are three major folds: a central,north-trending anticline; a shallow syncline on the west;
and a deep syncline on the east. The origin of these folds is
not clear, but they may be due to compression in the hingeof the east-tilted Sandia fault block during the lateCenozoic. Alternatively, the folds may be the result of theconvergence of the Tijeras and Gutierrez faults at thesouthern end of the Tijeras graben.
Road log: Day 1, Stops 2-5
Return to the vans and proceed east on 1-40 to the Cedar
Crest exit. Meshed and shot-creted outcrops on the left are
Pennsylvanian Madera Formation limestone and Permian Abo
Formation sandstone and shale that have been faulted andfolded along the Tijeras fault. The Ideal Cement Co. plant andquarry on the dip-slope to the south exploit Madera Formation
limestone.
Proceed north on NM-14. At 1.7 mi from 1-40 on eitherside of the road are hogbacks of Mesa Verde Group sandstone on the northwestern limb of the Tijeras syncline. Continue through Cedar Crest to the parking lot of the Mountain
Christian Church, on the east side of NM-14, 4.7 mi from 140.
Stop 2. Mountain Christian Church exposure of Jurassic
Todilto Formation gypsum. Gypsum was mined here by
Ideal Cement Co. and has been recently used by sculptors,
among whom it is known as Cedar Crest alabaster. The
Todilto Formation consists of about 3 m of thinly bedded,
fetid, lacustrine limestone overlain by 0 to 30 m of gypsum.
The gypsum contains carbonaceous matter which gives it amarbled appearance. On Day 2, the field trip will visit theIron Vein prospect in the Ortiz Mountains, where faulting

and hydrothermal activity have considerably altered the appearance of the Todilto gypsum.
Leave Mountain Christian Church, heading north on NM14. Immediately after leaving, note the Entrada Sandstone(Jurassic) and Chinle Formation maroon mudstones (Triassic) in the right roadcut. Continue 12.7 mi to Rt. 344. At
about 6 mi, note the Monte Largo Hills, a horst-block ofPrecambrian gneiss and granitic rocks on the right. Low
hills to the left at about 12 mi are composed of Precambrianquartzite and schist. San Pedro Mountains are directly infront. Continue on NM-14 about 1 mi to Rt. 344, turn right,
go 1.9 mi and turn left on the dirt road leading to the San
Pedro mine and mill.
Stop 3. San Pedro mine and mill. Leached outcrops ofcopper-skarn ore may have been worked here in Spanish
colonial times. Intermittent activity occurred during the
19th and the first half of the 20th centuries (Atkinson, 1961;
Elston, 1967). Goldfield Mining operated the San Pedro
mine in the 1960's and 1970's.
The skarn at San Pedro is developed in limestone of theMadera Formation. The intrusive body responsible for metasomatism was believed to be a monzonite-porphyry dike
about 300 m north of the mine (Lindgren et al., 1910;
Atkinson, 1961). The MoS 2 -bearing San Lazarus stock is themore likely source of the mineralization at the San Pedro mine
and at the pipe-shaped Carnahan silver—lead—zinc replacement deposit 1000 m south of the San Pedro mine (T.
Thompson, pers. comm. 1990).
Return to vans and drive back to NM-14. Head north on
NM-14. Mountain group to the north is the Ortiz Mountains.
One mile from the Rt. 344 junction enter the hamlet ofGolden, scene of a considerable gold rush in the 1830's, thefirst such boom west of the Mississippi. Gold was produced
from dry placers, largely through "coyoteing" or sinking ashaft through the poorly consolidated gravels and following
channels on the bedrock surface.
Continue north out of Golden on NM-14 about 7 mi toStagecoach Pass. The road in Stagecoach Canyon, on thesouth side of the pass, is cut into Upper Cretaceous Mancos
Shale which has been invaded by mainly concordant massesof Oligocene latite—andesite porphyry (Fig. 4).
Stop 4. Stagecoach Pass. Sills of Oligocene latiteandesite
porphyry intruding Upper Cretaceous Mancos Shale.
Concordant bodies of latite—andesite porphyry (dated at34.0 ± 2.2 Ma by Bachman and Mehnert (1978)) represent
the first phase of intrusion during the Oligocene. Johnson

62(1904), Lindgren et al. (1910), and Stearns (1953) interpreted the San Pedro—Ortiz porphyry belt as being dominated by laccoliths. While this conclusion seems correct in
the Cerrillos Hills (Disbrow and Stoll, 1957) recent investigations (Maynard et al., 1990) in the Ortiz Mountains havefailed to identify a laccolithic source of the sills found in therange. The sills are thicker on the western side of the Ortiz
Mountains, suggesting that the source laccolith was west of
the present Ortiz Mountains and has been eroded.
Return to vans and go north on NM-14. At 1.8 mi from
Stagecoach Pass come to commanding view of Cerrillos
Hills, Santa Fe, and the Sangre de Cristo Range. Descend
into Madrid. Just before entering Madrid, note exposures ofunmetamorphosed Menefee shales. The Menefee Formation,
composed of continental shale and fluvial sandstone, is hostto the coal that was mined by the Albuquerque and Cerrillos
Coal Co. from about 1900 until 1958 (Beaumont, 1979). On Day 2 the field trip will visit outcrops ofmetamorphosed Menefee shales in the Ortiz Mountains.
Continue north through the ramshackle of Madrid about
5 mi to the turnoff to the bucolic village of Cerrillos. Cerrillos has served as a set for motion pictures in recent
years, among them the western Young Guns. Drive into
Cerrillos, turn right at the first cross street, cross therailroad tracks, go straight, bear left at the cattle guard,
cross San Marcos Arroyo, then bear right, continuing 1.3
mi up a narrow arroyo to rip-rap quarry on left. This isprivate land, belonging to Cerrillos Gravel Products. Be
careful of large gravel trucks. Mineral rights are presentlyleased by Placer Dome Mining.
Stop 5. Cerrillos Sand and Gravel Co. rip-rap quarry.
The walking tour across the Cerrillos porphyry copper deposit to the Mt. Chalchihuitl turquoise mines starts here.

Geologic outline of the Cerrillos mining district, Cerrillos porphyry copper deposit,
and associated mineralization
David L. Giles
The 80 km2 Cerrillos mining district is north of the community of Cerrillos and about 25 km southwest of Santa Fe in
Santa Fe County, New Mexico. The area has a rich history ofturquoise, silver, and lead mining dating back over 1000
years. Turquoise deposits at Mt. Chalchihuitl (Figs. 5, 6) aresaid to be "the most extensive prehistoric mining operationsin North America" (Pogue, 1915) and date to 9501150 A.D.
(Warren and Mathien, 1985). Artifacts from the Mina del
Tiro and nearby Bethsheba lead—zinc—silver mines indicatethat Indians were also mining copper oxides for pigment andlead for pottery glaze at about the same time. Under Spanish
Colonial rule in the 16th and 17th centuries, it is rumored thatIndians were conscripted to work the mines, and that adisastrous cave-in in one of them was the immediate cause of
the Pueblo Revolt of 1680 (Schroeder, 1979). There is,however, no substantive historical evidence of this (Goodman
and Levine, 1990).
Some metal mines were worked intermittently by the
Spanish for lead, copper, and silver when they returned after
the revolt, and again under Mexican rule from 1821 until
1845. During the Territorial period, discovery of new leadzinc—silver lodes in the district in 1879 led to a boom thatlasted several years before subsiding to sporadic production
from a handful of mines. Operations had effectively ceasedby the mid-1950's. Elston (1967) reports production from1909 until 1957 as: 26,816 tons of ore; 734+ oz Au; 29,864
oz Ag; 181,224 lb Cu; 1,578,340 lb Pb; and 1,935,727 lb
Zn. Most of this recorded production came from the Tom
Payne and Cash Entry mines (Fig. 5).
Total production of turquoise from the known mines andmany small workings and pits throughout the district is
unknown. Early Indians of the region placed great value on
turquoise as a trade, ornamental, and ceremonial item. They
mined the Cerrillos Hills deposits, specifically Mt. Chalchihuitl, continuously for nearly 700 years to around 1700,
and thereafter intermittently to the present (Warren and Mathien, 1985). The Spanish, on the other hand, had no use forthe mineral.
Northrop (1959) reports that at least $2,000,000 worthof high-grade gem material was extracted from Turquoise
Hill (also known as the Tiffany mine) in the 1890's (Fig.
5). In 1959, Bear Creek Mining Company, a subsidiary of
Kennecott Copper, discovered and explored a small, lowgrade porphyry copper deposit in the southern part of the
district (Wargo, 1964). This deposit was then extensively
studied by Occidental Minerals Corporation in the mid1970's for possible in-situ leaching of the copper-oxide
portion of the deposit (Akright, 1979).
The Cerrillos Hills are at the northern end of the San
Pedro—Ortiz porphyry belt, a north-trending group of monzonitic stocks, laccoliths, sills, and dikes of Oligocene age(Woodward and Ingersoll, 1979; Maynard, this volume). The
Cerrillos intrusive/extrusive complex lies in a tilted structural
block bounded on the west by the east margin of Rio Grande
rift (Rosario and La Bajada faults), and on the south and eastby the northeast-trending Tijeras—Canoncito fault system(Lisenbee et al., 1979; Woodward, this volume). The
Paleozoic to Tertiary strata of the block dip regionally
eastward at 10-15° into the Santa Fe embayment at thesouthern end of the Espanola basin (Kelley, 1978). The
Cerrillos block comprises the "La Bajada constriction"(Kelley, 1952), a positive salient separating the rift graben of
Espanola basin from that of the Albuquerque basin to thesouth. No major mineral-controlling fault system similar to the
Tijeras—Canoncito has been recognized in the Cerrillos
district. Veins occupy northeast-trending fracture swarms inthe southern and the northern parts of the area (Figs. 5, 6).
At least four major periods of Oligocene igneous activity,
at a variety of separate or overlapping intrusive centers,
have been recognized in the Cerrillos Hills (Disbrow andStoll, 1957; Bachman and Mehnert, 1978). The rocks areuniversally quartz-poor and range from syenite to diorite,with monzonite predominating. The intrusive complex consists of stocks, plugs, sills, laccoliths, and dikes containing awide variety of mineral compositions and textures. Exposed
strata invaded by the complex include the Morrison
Formation (Jurassic); Dakota Sandstone, Mancos Shale, andMesaverde Group (Upper Cretaceous); Galisteo Formation

(Eocene); and Espinaso Volcanics (Oligocene) (Fig. 5). The
latter consists of a series of tuffs, agglomerates, and latite
flows that erupted contemporaneously with various intrusive
episodes throughout the Cerrillos Hills and the Ortiz Mountains (Kautz et al., 1981; Smith et al., this volume). These
stratigraphic units have been upturned along the margins of
the intrusive complex and occur as scattered inliers within.
The porphyry copper deposit and associated mineralizationin the southern Cerrillos district is centered on a sub-circularmultiple intrusive complex about 4 km in diameter (Fig. 6),
herein termed the San Marcos intrusive complex (SMIC).
The SMIC lies directly north of Cerrillos village and isisolated from other intrusive centers in the district by a skirt
of invaded wallrock, mainly Mancos Shale (Niobrara
Member) and Espinaso Volcanics (Fig. 5) (Bachman, 1975).
The nature and sequence of igneous events in the SMIC areunclear due to variable lithologies and hydro-thermal
overprint on the rocks. For simplicity, the sequence isgrouped here as follows: monzonite porphyry (oldest)
forming the core of SMIC; syenomonzonite as a small ameboid stock (about 1 km2) in the center of SMIC; and an
equigranular monzonite as a small, circular plug about 360
m in diameter, emplaced into a lobe of the syenomonzonite
mass (Fig. 6). Porphyry copper mineralization is spatially,
and in part genetically, related to the young monzonite plug.
The monzonite porphyry is dark to medium gray andcharacterized by hornblende and plagioclase phenocrysts
ranging in size and abundance. Most of the monzoniteporphyry sills and laccoliths were emplaced into the relatively incompetent Mancos Shale. About 1.5 km up theaccess road to the copper deposit from San Marcos Wash,"fingers" of monzonite appear to be gently probing into theshales; possibly they are the leading edge of the sill-likemass on the opposite side of the arroyo. Contact effects inthe shale are limited to thin baked selvages and deformation
is minimal.
The syenomonzonite is pinkish gray, equigranular tocoarsely porphyritic with augite (or biotite) and K-feldspar
phenocrysts, and is cut by numerous pink aplite dikes. The
late monzonite plug is light gray, coarse-grained equigranular, and forms a barren core around which a stockwork
fracture zone has localized a ring of copper mineralization(Fig. 6). Scattered lesser rocks in the area include pre- andpost-mineral monzonite dikes, late basalt(?) dikes, and flows
probably belonging to the Espinaso Volcanics.
The Cerrillos copper deposit displays most of the alteration and mineral zoning features typical of hydrothermal
porphyry copper systems. It is a low-grade (about 0.3% Cu),
MoS 2 -poor, low total-sulfide system with a high chalcopyrite/pyrite ratio in a zone coincident with strong stockwork fracturing, local vein and replacement silicification,and potassic alteration (secondary biotite and K-feldspar).
There is no supergene chalcocite enrichment blanket at
depth (Wargo, 1964; Akright, 1979), although secondary
chalcocite did form at the present exposure level from earlier
oxidation and leaching. The deposit has a well-developed
pyritic halo characterized by widespread propylitization(chlorite and epidote), local clay-sericite alteration, andnear-background copper values.
The copper ring, as defined by the >2000 ppm Cu-inrock
surface contour (Fig. 6), ranges from 30 to 180 m wide andencircles two-thirds of the monzonite plug. Drillingindicates that the deposit is a near-vertical cylinder adjacentto the walls of the barren monzonite plug. From the sharpinner boundary (an assay wall), copper grades decline gradually outward to an economic cutoff in the pyrite zone. The
bulk of the copper deposit overlaps, for the most part, theporphyritic syenomonzonite lobe and its contact with theolder monzonite porphyry (Fig. 6). The above relationships
suggest that intrusion of the late monzonite plug was instrumental in ground preparation (stockwork fracturing)around its border, thus creating a zone of permeability forlater copper-bearing hydrothermal solutions. This is the primary control of mineral deposition.
The primary sulfides in the Cerrillos deposit are chalcopyrite and pyrite, with trace amounts of bornite, marcasite,
enargite, and molybdenite. Minor amounts of secondary
chalcocite and covellite occur as relicts in the oxide zone
and as supergene minerals in the mixed oxide—sulfide zone
at depth. Significant magnetite occurs as veinlets both preceding and crosscutting primary copper mineralization, andas breccia cement. Small amounts of gold (in electrum) arepresent. The gold seems to correlate with zones of stronger
silicification, potassic alteration, higher copper values, andmagnetite veining. This genetic association has been discussed by Sillitoe (1979) and Cox and Singer (1988). All
mineralization in the deposit is strongly fracture-controlled;
disseminated oxides and sulfides are rare.
Late lead—zinc veins and veinlets crosscut copper mineralization and occur in all rock types in the deposit areas,
including the barren monzonite core. Extensive surface rockchip sampling shows no halo-type distribution for lead andzinc peripheral to copper; in fact, many lead—zinc highs
coincide with copper highs. It would seem that lead and zinc
did not participate in true spatial zoning, as in manyother porphyry systems, but were superimposed on the copper
deposit along a swarm of northeast-trending veins from Minadel Tiro to Cash Entry. Similarly, the surface distribution ofmolybdenum does not mimic that of copper, as might be
expected, but is erratic and equally prominent in both thecopper zone and pyrite halo.
The principal secondary copper minerals in the oxide cap
are neotocite and brochantite, with lesser chrysocolla andpitch limonite, and small amounts of malachite, azurite,
cuprite, antlerite, melaconite (black earthy tenorite, CuO),
hematite, residual chalcocite, and native copper. These minerals formed in a supergene environment where the copper/
sulfur ratio was very high, and where the copper-sulfide suite
prior to most recent oxidation was a mix of primarychalcopyrite and local secondary chalcocite. Two areas ofbetter grade and potentially minable oxide copper have beendelineated by detailed drilling in the southwest and southeast
quadrants of the deposits. The goethitic-oxide cap ranges
from 30 to 120 m thick, with a mixed sulfide—oxide zone at
the interface and primary sulfides at depth. The variation is
attributed by Akright (1979) to post-oxidation faulting.
The makeup of the limonite and copper-oxide minerals in
the outcropping cap allows certain observations about thenature of the eroded higher levels of the deposits as well as
the remaining deposit at depth:
1) upper parts of the deposit (now removed) had a high pyrite
and high total-sulfide content, resulting in leaching ofcopper and precipitation of secondary chalcocite at about
the present exposure level;2) the precursor sulfide deposit at the present level was low
in total sulfides (about 1-2%) and this sulfide suite had

65
more than 50% chalcopyrite and, locally, more than
one-third chalcocite;3) very little copper has been leached from the cap at thepresent exposure level;4) the probability of a significant second-generation chalcocite blanket at depth is nil;5) surface copper grades can be predicted to be essentiallythe same as in the sulfide zone at depth, as confirmedby drilling (Wargo, 1964; Akright, 1979; unpublisheddata).
The pyrite halo surrounding the copper deposit is subcircularand about 1.7 km in diameter. The halo averages about 2%
pyrite, but contains scattered patches of bleached and argillized rock that initially contained about 8 to 10% pyrite.
These bleached zones are characterized by jarositic limonite
and numerous small occurrences of turquoise (Fig. 6) withsporadic variscite (AlPO 4 .2H 2 0).
The turquoise shows have been prospected off and on by
locals for many years, with occasional discovery of gemquality material. An important prehistoric turquoise deposit
occurs at Mt. Chalchihuitl about 1.2 km north of the equipment yard (Fig. 6). The name of the mine is a derivative of
the word for turquoise in Nahuatl, the Aztec dialect, thus
imparting a sense of its age. This mine is in clay-altered andweakly tourmalinized volcanics, probably in inlier of
Espinaso Volcanics. A nearby drill hole collared in the unit
penetrated monzonite porphyry at about 30 m. The turquoise

occurs as nodules in open spaces and as irregular veinlets
with little continuity. Turquoise and variscite are secondaryoxide minerals that require the presence of free phosphorus
in the supergene environment in order to form. The logicalsource of phosphorus is apatite, which is reported as a relatively abundant accessory mineral in the monzonites. Presumably, supergene acid conditions in the high pyrite zones
were strong enough to break down apatite and release
phosphorus to ground waters.
Several lead—zinc—silver vein deposits in the southern
Cerrillos district have had significant production. Most notable are the Cash Entry mine and Mina del Tiro (Fig. 6).
According to Spanish Archives (Goodman and Levine,
1990), the latter was mined by the Spanish to obtain lead formusket balls. In general, the vein ores consist of galena andsphalerite, with lesser chalcopyrite and pyrite, in a gangue ofankerite and quartz. A little barite, stibnite, bournonite,
wulfenite, rhodochrosite, amethyst, and chalcedony are reported (Disbrow and Stoll, 1957). Vein material occurs as
lenses, streaks, crustifications, and disseminations in shootsthat partially fill northerly trending shear zones; wallrock ismostly sericitized monzonite. Silver is associated with galena, and gold probably with chalcopyrite. Historically, average ore from the district yielded only 1.1 oz Ag and 0.03
oz Au per ton, but some small high-grade shipments contained over 600 oz/ton Ag and up to 5 oz/ton Au.

Geology and gold mineralization of the Ortiz Mountains
Stephen R. Maynard
Introduction
Approximately 350,000 oz (11,000 kg) of gold havebeen produced from the Ortiz Mountains since the 1820's.
The 260 km2 Ortiz Mine Grant was secured from the
Mexican government in 1833 by Jose Francisco Ortiz forthe purpose of mining gold. Surface rights were separated
from mineral rights in the 1940's and are divided among
several owners. The mineral rights on the Grant haveremained more or less intact to the present. Mineral rights
are currently owned by Ortiz Mines, Inc. of Joplin,
Missouri, and are leased to Pegasus Gold, which is operatorof an exploration, development, and production joint
venture with LAC Minerals, USA, Inc.
Igneous rocks
Igneous rocks in the Ortiz Mountains (Fig. 7) can bedivided into an earlier calc-alkaline group and a later
alkaline group. Isotopic-age determinations on these groups
constrain mid-Tertiary movement on the Tijeras—
Canoncito fault system (TCFS) to the period 34 to 30 Ma.
Most gold mineralization is spatially and temporally related
to the alkaline igneous rocks.
Latite—andesite porphyry sills, laccoliths, and dikes and agranodiorite stock comprise the calc-alkaline group. The K—
Ar age determination that best fits with geologic relations on
the latite—andesite porphyry is 34.0 ± 2.2 Ma (Bachman andMehnert, 1978). This date was determined from hornblende in
unaltered rocks on the west side of the Ortiz

Mountains. Older age determinations of 43.2 ± 2.3 Ma (Kay,
1986) and 47.1 ± 3.2 Ma (Bachman and Mehnert, 1978)
appear to contradict geologic relations and may represent
contamination by upper crustal material. Latite—andesite
sills appear to have propagated from laccolithic masses on
the west side of the Ortiz Mountains and are cut by the
TCFS. The sills generally intrude less competent lithologies,
particularly shales and mudstones of the Menefee, Mancos,
Morrison, and Chinle Formations, and show little evidence ofthermal metamorphism at the contacts. Latite—andesite sills
are important brecciated hosts for gold mineralization in the
Carache Canyon breccia-pipe area.
The Candelaria granodiorite stock is probably the youngest
phase of calc-alkaline magmatism, although direct crosscutting evidence with the latite—andesite sills is not available.
A large hornfels aureole is present around the stock. The
garnet—pyroxene skarn-altered Greenhorn Limestone at Lukas Canyon is within this contact aureole and the Candelaria
granodiorite is presumed to be the thermal source for both thehornfels aureole and the skarn.
Alkaline rocks are represented by a large nepheline-bearing augite-monzonite stock, a latite-porphyry hypabyssal
plug, the Ortiz diatreme (vent breccia), and radial trachyticlatite dikes. K—Ar age determinations range from 30 to 26
Ma for the alkaline rocks (Kay, 1986). The 29.6 ± 1.5 Ma
augite-monzonite stock intruded the Eocene Galisteo Formation to within 1000 m of its projected surface, cut acrossthe Ortiz graben (the principal structure of the TCFS in theOrtiz Mountains), and produced a wide contact-metamorphic aureole in the surrounding sedimentary rocks. Thisaureole is defined by intense hornfels development and iseasily mapped by topographic highs underlain by what
would otherwise be recessive sedimentary rocks. The Ortiz
diatreme and its related hypabyssal latite-porphyry plug cut
the augite-monzonite stock. Dikes radial to the diatreme
crop out in the distance of approximately 3.5 km.
Structural geology
The Tijeras—Canoncito fault system dominates the structural geology of the Ortiz Mountains (Figs. 7, 8). The Ortiz
graben, the Ortiz diatreme, and the Carache Canyon breccia
pipe appear to be structurally controlled by the TCFS.
The Ortiz graben contains several tilted blocks ofGalisteo and Menefee Formations. Vertical stratigraphic
separations of 600-1000 m are estimated on the graben'sbounding faults. Left-lateral stratigraphic separation of the
Dakota Formation measures approximately 5 km across the
TCFS. The northern side of the TCFS is downthrown withrespect to the southern side. The true amount and directionof displacement are unknown.
The Ortiz diatreme, or vent breccia, erupted on the northwestern side of the northwestern graben-bounding fault andpartly occupies it. The diatreme forms a crude ellipse approximately 2100 by 1200 m in plan. The diatreme has been
drilled to a depth of 500 m (Fig. 9). It is composed principallyof tuff, lithic tuff, and volcanic breccia. Lithic clasts in the tuff
are composed of augite monzonite, latite—andesite

porphyry, and Cretaceous and Eocene sedimentary rocks.
Breccia pipes occur in two principal zones on the southeastern and northwestern margins of the diatreme. The
northwestern margin is a fault contact of vent breccia withaugite monzonite. Substantial throw is suspected on thisfault, bringing the vent breccia down with respect to theaugite monzonite. The southeastern margin is an intrusive
contact characterized by intense brecciation of the wallrocks as at the Cunningham Hill and Benton mines(Lindqvist, 1980; Wright, 1983; Kay, 1986).
The southeastern margin of the Carache Canyon collapsebreccia pipe lies about 120 m to the northwest of the north-western margin of the Ortiz graben (Fig. 10). The pipe isapproximately 625 by 312 m, elongate in a northeast—southwest direction, and plunges at 75° toward the southwest.
It is hypothesized that deep-seated fractures parallel to
the TCFS served as a conduit for vapor release and as azone of weakness for brecciation and subsequent collapse,
forming the Carache Canyon breccia pipe. The Ortiz diatreme formed 2000 m to the northest, parallel to the strikeof the Golden fault. Magmatic withdrawal during eruptionof the Ortiz diatreme may have allowed room for the collapse of the Carache breccia pipe.
Gold mineralization
Lode-gold deposits in the Ortiz Mountains generally occur along or near strands of the TCFS. Lodes are ofvarious types including breccias, veins, skarns, andporphyry-type occurrences. A variety of breccias host gold
mineralization in the Ortiz Mountains. Placer-gold
deposits in Plio-Pleistocene gravels lie on the eastern andsouthern flanks of the mountains.
Road log: Day 2
Retrace Day 1 route north on NM-14 to Gold Mine Road,
2.8 mi north of Madrid. Note outcrops of Eocene Galisteo
Formation sandstone in gulch north of Madrid. Turn south
on Gold Mine Road, bear right at 2.3 mi and continue 4.3
mi to the gate and office of the Cunningham Hill mine. Go
through gate and drive about 0.2 mi to vista to northeast.
Stop 1. Overview and discussion of Tijeras—Canoncito
fault system northeast of the Ortiz Mountains.

Return to the vans and head west on the road past themaintenance building and tailings piles on the left and theleach pad, mill site, and primary and secondary crusher
locations on the right. Park at the east end of the Cunningham
Hill open pit.
Stop 2. Cunningham Hill mine. Micron-size gold is commonly associated with crystals of pyrite, magnetite, calcite,siderite, quartz, chalcopyrite, or scheelite in open spaces
between clasts of brecciated Galisteo Formation at Cunningham Hill. Kay (1986) stated that the Cunningham Hill
breccia is least permeable and lowest grade on the southernand eastern borders, and most permeable and highest grade on
the northern and western sides bordering the vent breccia(Figs. 8, 9, 11). Kay (1986) postulated that the ore-hosting
breccia formed later than the diatreme due to explosive vapor
release and he attributed the downward tapering of the breccia
body to a primary stratigraphic control. That is, incompetentshales below the ore-hosting sandstone body could not
maintain the open spaces.
Consolidated Gold Fields extracted approximately 250,000oz (7800 kg) of gold from the Cunningham Hill open-pit minein the period 1979-1987. Mined ore averaged 0.055 opt in 6
million tons (1.9 gmt in 5.4 million metric tons).
Board the vans on the western side of the pit at thescheelite stockpile and head west 0.9 mi past the old Ortiz
mine. The Ortiz mine is the site of the original lode discovery
in the Ortiz Mountains in 1932 and is the geographic center ofthe 260 km2 Ortiz Mine Grant. The Ortiz, English,

Shoshone, and Florencio mines lie on, or near, a N20Etrending fault contact between the augite-monzonite stock andthe Ortiz diatreme (Fig. 9). Gold appears to be associatedwith magnetite fillings in a wide brecciated zone in the augite
monzonite at the Ortiz and English mines. At the Florencio
mine, the brecciated host is quartzite developed in Mesa
Verde or Galisteo sandstone.
Continue in vans up the Cunningham Gulch road as far as
the vans can go, then walk through the Cunningham Gulch
porphyry copper—gold prospect.
Stop 4. Cunningham Gulch porphyry copper—gold prospect. The hypabyssal latite plug in Cunningham Gulch appears to be the barren core of a low-grade gold-bearing
copper porphyry system (Fig. 8). Crackle breccia and stockwork veinlets filled with tenorite + magnetite ± fluorite
molybdenite are associated with gold mineralization.
Included in this stop are outcrops of the barren porphyry coreand roadcuts showing the copper—gold mineralization.
Continue on foot along the Cunningham Gulch road overa pass, then down to the Carache Canyon decline portal
where the field trip will meet the vans for lunch. Just before
arriving at the pass, the road crosses the Golden fault, thenorthwest-bounding fault of the Ortiz graben. Rocks along
the road down to the decline portal are quartzite derived
from the Galisteo Formation. After lunch, vans will takethe group up Carache Canyon for a walking tour of the
Carache Canyon breccia pipe.
Stop 5. Carache Canyon. The Carache Canyon deposit isassociated with a collapse-breccia pipe measuring 550 by
300 m on the surface, plunging 70° S4OW, and having aknown down-plunge extent of over 1000 m (Fig. 10). The
pipe is characterized by collapsed beds of shale and sandstone and latite—andesite porphyry sills. A relict
stratigraphy, largely defined by the distinctive Point Lookout
Sandstone, shows as much as 250-300 m of collapse. The
breccia typically is clast-supported and locally contains asmuch as 10% dark-colored rock-flour matrix.
Secondary fracturing of the Carache Canyon pipe resulted
in randomly oriented, open-space fractures that are concentrated in an annular zone around the southwestern margin of
the pipe. The secondary fracture zones occur inside andoutside the pipe margins. The open-space fracture zones
developed best in the latite—andesite sills and in the Point
Lookout Sandstone. In contrast, the shales of the Mancos
and Menefee Formations contain few open-space fractures.
Thus, outside the pipe, mineralization is concentrated in aseries of stacked, tabular bodies separated by intervals of
relatively barren shale.
Unoxidized mineralization at Carache Canyon is in theform of coarse gold in open-space fractures with pyrite,
pyrrhotite, chalcopyrite, sphalerite, galena, arsenopyrite,
calcite, and adularia. Gold appears to have been depositedlate in the paragenetic sequence (Coles, 1990). Gold mineralization at Carache Canyon appears to be closely related
temporally to the intrusion of trachytic-latite dikes. Somedikes are fractured and mineralized. Others cut mineralization and are unfractured.
In May 1988, LAC Minerals announced the discovery of ageologic resource of 5 million tons grading 0.094 opt (4.5
million metric tons of 2.9 gmt) gold at Carache Canyon.
Return to the vans and drive out of Carache Canyon about 2
mi to the Lone Mountain Ranch road. The Ortiz surface, theregular plain sloping away from the Ortiz Mountains, isunderlain by the late Tertiary Tuerto Gravel, host of thenumerous placer deposits in the area. Turn right, cross thecattle guard, take the next right and proceed approximately 2
mi to the Iron Vein prospect.
Stop 6. Iron Vein prospect. A tectonized solution-collapse
breccia is developed in the gypsum member of the Jurassic
Todilto Formation along the southeastern margin of the
Ortiz graben. The solution-collapse breccia is apparently aregional feature of probable Tertiary age. At the Iron Vein,
the breccia consists of clasts of Jurassic Morrison Formation
sandstone and shale supported by a yellowish-orange matrixof sand, calcite, clay, and gypsum. In addition, magnetite +hematite ± garnet skarn is locally developed in the Todilto
limestone. Controls on gold mineralization in the Iron Vein
area appear to be related to faulting and to an underlying
granodiorite stock.
Return to vans, drive back to Lone Mountain Ranch road,turn right and continue 0.9 mi to power-line access road.
Follow power-line road 0.9 mi, turn right. Follow road backto north, pass windmill at 0.6 mi, go 0.1 mi, turn right, and go
0.6 mi to entrance to Lukas Canyon.
Stop 7. Lukas Canyon. The Lukas Canyon deposit, acopper—gold, garnet—pyroxene skarn developed in theGreenhorn Limestone Member of the Mancos Formation,
contains an announced resource of 0.03 opt gold and 0.25%
copper in 6 million tons (1.02 gmt gold in 5.4 million metrictons). Gold and copper mineralization postdates the primaryskarn and is believed to be related to a late retrograde alteration event when the garnet—pyroxene skarn was altered69

to chlorite-actinolite-epidote. Sulfides are locally present as
pyrite-chalcopyrite but the deposit has been intensely
oxidized, resulting in a complex assemblage of copper- andiron-oxides. The skarn crops out on a dip-slope, making it
amenable to surface mining.
End of excursion.

Combined references
Akright, R. L., 1979, Geology and mineralogy of the Cerrillos copper
deposit, Santa Fe County, New Mexico: New Mexico Geological Society, Guidebook 30, pp. 257-260.
Atkinson, W. W., 1961, Geology of the San Pedro Mountains, Santa Fe
and Sandoval Counties, New Mexico: New Mexico Bureau of Mines &Mineral Resources, Bulletin 77, 49 pp.
Bachman, G. 0., 1975, Geologic map of the Madrid 15' quadrangle, Santa
Fe and Sandoval Counties, New Mexico: U.S. Geological Survey, Geologic Quadrangle Map GQ-l268.
Bachman, G. 0., and Mehnert, H. H., 1978, New K-Ar dates and the late
Pliocene to Holocene geomorphic history of the central Rio Grande
region, New Mexico: Geological Society of America, Bulletin, v. 89,
pp. 283-292.
Beaumont, E. C., 1979, Geology of the Cerrillos coal field, Santa Fe
County, New Mexico: New Mexico Geological Society, Guidebook 30,
pp. 269-275.
Black, B. A., 1979, Structure and stratigraphy of the Hagan
embayment: a new look: New Mexico Geological Society,
Guidebook 30, pp. 101105.
Black, B. A., and Hiss, W. L., 1974, Structure and stratigraphy in thevicinity of the Shell Oil Co. Santa Fe Pacific no. 1 test well, southern
Sandoval County, New Mexico: New Mexico Geological Society,
Guidebook 25, pp. 365-370.
Bruns, J. J., 1959, Petrology of the Tijeras Greenstone, Bernalillo
County, New Mexico: Unpublished MS thesis, University of New
Mexico, Albuquerque, 119 pp.
Coles, D., 1990, Alteration and mineralization of the Carache Canyon
breccia pipe, Santa Fe County, New Mexico: Unpublished MS thesis,
Colorado State University, Fort Collins.
Connolly, J. R., 1981, Geology of the Precambrian rocks of Tijeras Canyon, Bemalillo County, New Mexico: Unpublished MS thesis, University of New Mexico, Albuquerque, 147 pp.
Connolly, J. R., 1982, Structure and metamorphism in the Precambrian
Cibola gneiss and Tijeras greenstone, Bemalillo County, New Mexico:
New Mexico Geological Society, Guidebook 33, pp. 197-202.
Cox, D. P., and Singer, D. A., 1988, Distribution of gold in porphyry
copper deposits: U.S. Geological Survey, Open-file Report 88-46,
23 pp.
Disbrow, A. E., and Stoll, W. C., 1957, Geology of the Cerrillos area,
Santa Fe County, New Mexico: New Mexico Bureau of Mines &Mineral Resources, Bulletin 48, 73 pp.
Elston, W. E., 1967, Summary of the mineral resources of Bernalillo,
Sandoval, and Santa Fe Counties, New Mexico: New Mexico Bureauof Mines & Mineral Resources, Bulletin 81, 81 pp.
Fulp, M. S., Cavin, W. J., Connolly, J. R., and Woodward, L. A., 1982,
Mineralization in Precambrian rocks in the Manzanita-north Manzano
Mountains, central New Mexico: New Mexico Geological Society,
Guidebook 33, pp. 303-304.
Goodman, L. J., and Levine, D. F., 1990, The mines of the Cerrillos
District, New Mexico; myths and realities: El Palacio (Museum of New
Mexico, Santa Fe), v. 96, no. 1, pp. 20-39.
Kautz, P. F., Ingersoll, R. V., Baldridge, W. S., Damon, P. E., andShafiqullah, M., 1981, Geology of the Espinaso Formation(Oligocene), north-central New Mexico, part 2: Geological Society ofAmerica, Bulletin, v. 92, pp. 2318-2400.
Johnson, D. W., 1904, Geology of the Cerrillos Hills, New Mexico: New
Era Printing Co., Lancaster, Pennsylvania.
Kay, B. D., 1986, Vein and breccia gold mineralization and associatedigneous rocks at the Ortiz mine, New Mexico, USA: Unpublished MS
thesis, Colorado School of Mines, Golden, 179 pp.
Kelley, V. C., 1952, Tectonics of the Rio Grande depression of central
New Mexico: New Mexico Geological Society, Guidebook 3, pp. 93105.
Kelley, V. C., 1977, Geology of Albuquerque basin, New Mexico: New
Mexico Bureau of Mines & Mineral Resources, Memoir 33, 60 pp.
Kelley, V. C., 1978, Geology of Espanola basin, New Mexico: New
Mexico Bureau of Mines & Mineral Resources, Geologic Map 48.
Kelley, V. C., and Northrop, S. A., 1975, Geology of Sandia Mountains
and vicinity, New Mexico: New Mexico Bureau of Mines & Mineral
Resources, Memoir 29, 136 pp.
Lindgren, W., Graton, L. C., Gordon, C. H., 1910, The ore deposits of
New Mexico: U.S. Geological Survey, Professional Paper 68, 361 pp.
Lindqvist, W. F., 1980, The exploration of the Ortiz gold deposit, New
Mexico-geology and exploration: Colorado Mining Association, 1980
Yearbook, pp. 106-112.
Lisenbee, A. L., 1975, Shale diapirism and the development of theGalisteo syncline; in Woodward, L. A., and Northrop, S. A. (eds.),
Tectonics and mineral resources of southwestern North America: New
Mexico Geological Society, Special Publication 6, pp. 88-94.
Lisenbee, A. L., Woodward, L. A., and Connolly, J. R., 1979, TijerasCanoncito fault system-a major zone of recurrent movement in northcentral New Mexico: New Mexico Geological Society, Guidebook 30,
pp. 89-99.
Lodewick, R. B., 1960, Geology and petrography of the Tijeras gneiss,
Bemalillo County, New Mexico: Unpublished MS thesis, University of
New Mexico, Albuquerque, 63 pp.
Maynard, S. R., Nelsen, C. J., Martin, K. W., and Schutz, J. L., 1990,
Geology and gold mineralization of the Ortiz Mountains, Santa Fe
County, New Mexico: Mining Engineering, v. 42, pp. 1007-1011.
Northrop, S. A., 1944, Minerals of New Mexico: University of New
Mexico Press, Albuquerque, 665 pp.
Picha, M. G, 1982, Structure and stratigraphy of the Montezuma SalientHagan Basin area, Sandoval County, New Mexico: Unpublished MS
thesis, University of New Mexico, Albuquerque, 248 pp.
Pogue, J. E., 1915, The turquoise; a study of its history, mineralogy,
geology, archaeology, mythology, folklore and technology: National
Academy of Sciences, Memoir 3, v. 12, pt. 2.
Schroeder, A. H., 1979, The Cerrillos mining area; in Ingersoll, R. V.
(ed.), Archaeology and history of Santa Fe country: New Mexico Geological Society, Special Publication 8, pp. 13-16.
Sillitoe, R. H., 1979, Some thoughts on gold-rich porphyry copper deposits: Mineralium Deposita, v. 14, pp. 161-174.
Stearns, C. E., 1953, Early Tertiary volcanism in the Galisteo-Tonquearea, north-central New Mexico: American Journal of Science, v. 251,
pp. 415-452.
Wargo, J. G., 1964. Geology of a primary disseminated copper deposit
near Cerrillos, New Mexico: Economic Geology, v. 59, pp. 1525-1538.
Warren, A. H., and Mathieu, F. J., 1985, Prehistoric and historic
turquoise mining in the Cerrillos district: time and place: Papers of theArchaeological Society of New Mexico, #10, 110 pp.
Woodward, L. A., and Ingersoll, R. V., 1979, Phanerozoic tectonic
setting of Santa Fe country: New Mexico Geological Society,
Guidebook 30, pp. 51-58.
Wright, A., 1983, The Ortiz gold deposit (Cunningham Hill)-geology andexploration: Nevada Bureau of Mines and Geology, Report 36, pp. 4251 .

71

New Mexico Bureau of Mines & Mineral Resources Bulletin 137, April 1991

Proterozoic tectonic history of the Manzano Mountains,
central New Mexico
Amy G. Thompson', Jeffrey A. Grambling', and R. David Dallmeyer²'Department of Geology, University of New Mexico, Albuquerque, New Mexico 87131; 'Department of Geology, University of Georgia, Athens, Georgia 30602

Introduction
Proterozoic rocks of central New Mexico crop out in anorth-trending, linear mountain chain along the eastern margin of the Cenozoic Rio Grande rift. The chain includes the
Sandia, Manzano, and Los Pinos Mountains. A major unconformity located near the top of each range separates thebasement rocks from overlying Paleozoic sedimentaryrocks. The unconformity varies from flat-lying to shallowly
east-dipping.
Proterozoic rocks below the unconformity record at least
two superimposed tectonometamorphic events. This study
describes these events within the central and southern Manzano Mountains. The study area consists of supracrustal rocks
metamorphosed to the amphibolite to greenschist grade,
together with two granitic plutons (Fig. 1). The supracrustal
rocks are divided into two distinct tectonometamorphic sequences, juxtaposed along a northeast-trending ductile shear
zone that is best exposed in Monte Largo Canyon, in thecentral Manzano Mountains. Kinematic indicators in both
tectonic sequences indicate that the southern, amphibolitegrade sequence was thrust to the northwest over the northern,
greenschist-grade sequence. A greenschist-grade metamorphic event overprinted both sequences during and after
thrusting; it resulted in prograde metamorphism of thegreenschist-grade crustal sequence and retrograde metamorphism of the amphibolite-grade sequence. Thrusting has
been dated at about 1350 Ma by 40Ar/39Ar thermochronology(Thompson et al., 1989).
The purpose of this field trip is to examine Proterozoic
structures and metamorphism within the Manzano Mountains. The focus will be on the style and timing of juxtaposition of the two tectonic sequences. The first day willconcentrate on the ductile shear zone, exposed in Monte
Largo Canyon, that separates the greenschist- and amphibolite-grade sequences. Metamorphic-grade and shearsense indicators in both tectonic sequences will be
examined. The second day will focus on rocks farther southin the southern tectonic sequence. The Priest pluton and theretrograded portion of the amphibolite-grade sequence inthe area near Abo Pass will be examined.
Highlights en-route to Stop 1
The field-trip route is shown in Fig. 2. From
Albuquerque, the trip travels south on 1-25. After 1-25
crosses the Rio Grande, the highway turns south throughroadcuts that expose late Pleistocene basalts overlying
unconsolidated sediments of the Santa Fe Group. The
Manzano Mountains are visible along the eastern horizon,
where the unconformity separating flat-lying Paleozoic
sedimentary units from underlying Precambrian igneousand metamorphic rocks can be seen.

On the right, west of the highway immediately past the
Los Lunas exit, is the Los Lunas volcano, one of many
volcanic cones that dot the Rio Grande rift. Its andesitic
lavas are dated at 1.3-1.1 Ma (Kudo, 1982 and references
therein). Directly to the east, across the Rio Grande valley,
the isolated, rounded knob is Tome Hill, another intrarift,
andesitic volcano; this one, however, is dated at 3.5 Ma(Kudo, 1982 and references therein).
The view east beyond Tome Hill to the piedmont near the
Manzano Mountain front shows, in the far distance, theHubbell Springs fault scarp. Based on a study of soils andfault-scarp degradation models, Machette (1982) reported alate Pleistocene to early Holocene age for the most recentintersection). Continue south for 5.5 mi and look for abrown sign on the right pointing to John F. Kennedy (JFK)
Campground. Turn left onto the dirt road at that sign.
Continue east on the dirt road. Approximately 5.6 mi eastof NM-47, a sign indicating a turn to the left toward JFKCampground is posted at a Y-intersection. Pass the signand continue east for an additional 0.2 mi to a barbed-wirefence. Here, the east-bound road intersects a north—southroad. Turn right onto the north—south road and travelsouth 2.2 mi to the first stop. The first stop is a view to thenortheast along strike of the Monte Largo shear zone. Also,
note the horizontal Paleozoic sediments capping the
Manzano Mountains to the north.
Stop 1: Overview of Monte Largo shear zone
The shear zone is a northeast-trending, southeast-dipping
ductile fault that emplaced amphibolite-grade rocks northwestward, on top of greenschist-grade rocks. The amphibolite-grade rocks southeast of the shear zone dip at asteeper angle than the greenschist-grade rocks to thenorthwest. This relation can be seen clearly in the distantmountain front (Fig. 3).

rupture on this fault. The scarp height ranges from 4 to 30
m. The Hubbell Springs fault is part of the Hubbell bench,an 88 km long fault zone considered to be the major fault
bounding the western edge of the Manzano Mountains(Kelley, 1982).
Continuing south on 1-25, use the first Belen exit (Business 1-25) to the second traffic light. Turn left (east) onto
NM-309 (Reinken Avenue). continue for 2.4 mi; pass overthe Rio Grande and approach a 4-way-stop intersection. At
that intersection, turn south (merge right) onto NM-47 (thefirst sign identifying NM-47 is about 0.25 mi south of the

Next, proceed to Stop 2 in Monte Largo Canyon where
the greenschist- and amphibolite-grade rocks and the tectonic contact that separates them will be examined. Retrack 2.2 mi north to the jog in the east-bound road, thenturn east. Pass a house with a large, white railroad-tank carin its front yard. Monte Largo Canyon is directly to theeast. Continue on the same road through the curve to thenorth and across the major arroyo that drains Monte Largo
Canyon, then turn right (east) on the first intersecting dirt
road. Continue on that road toward the mountain front,back across the arroyo; after crossing the arroyo, turn left
on the third dirt road. This road crosses the main arroyo
again. About 30 ft past the north side of the arroyo, anotherdirt road forks to the right and leads directly east into
Monte Largo Canyon. This road is fairly rough andrequires a vehicle with high road clearance.
Stop 2: Monte Largo Canyon shear zone
The objectives of the second stop are to examine themetamorphic grade of each tectonic sequence and to lookin detail at structures along the contact between the two
tectonic sequences. The field-trip route into Monte Largo
Canyon is shown in Fig. 4. The hike will start on the road

into the canyon and proceed approximately 800 m togreenschist-grade phyllites exposed north of the main creek(stop a). The hike will then continue 200 m up the creek
bed, then up a ridge to the south where its path will cross theshear zone in order to examine mylonitic fabrics in detail(stop b). Finally, the hike will proceed about 800 m farther
up-canyon where the route proceeds up a knob (to the east)
into hornblende—plagioclase mafic rocks of theamphibolitegrade sequence (stop c). After returning fromthe hike, retrace the route back to Belen. The trip willcontinue on a second day beginning with Stop 3.
Highlights en-route to Stop 3
The main features viewed on the second day are thecontacts of the Priest pluton and the overprinting relationshipbetween greenschist-grade and amphibolite-grade metamorphism, both best seen in the southernmost Manzano
Mountains. The trip starts in Belen. Follow the route for the
first day, again turning right (south) onto highway NM-47.
Drive for about 14 mi to a railroad crossing. Cross the tracks
and then turn left immediately onto the railroad access roadsouth of the tracks. Drive for 4.7 mi, paralleling the railroad
tracks. On the drive toward the mountain range, the Priest
pluton can be distinguished from supracrustal rocks becauseof differences in weathering, exfoliation, and its soil that
supports growth of abundant low bushes. The road trends
south for the last 0.3 mi to the railroad cut just north of Abo
Canyon, where the hike begins for Stop 3.
The view to the south is of the Los Pinos Mountains, thesouthern extension of the Manzano Mountains. The Los
Pinos granitic pluton, which intruded metamorphic rocks
similar to those in the southern tectonic sequence of the
Manzano Mountains, forms the prominent cliffs to the south.
Stop 3: Southern margin of the Priest pluton
The purpose of this stop is to examine evidence for agreenschist-grade, retrograde metamorphic event that overprints amphibolite-grade rocks, and to examine the contactof the Priest pluton. The field-trip route into the Abo Pass
area is shown in Fig. 5. The hike starts at the railroad cut
and proceeds upslope to the east and north through metasedimentary units for about 400 m (stop a). The first stop is
at a megascopic fold defined by quartzite and phyllite, thelatter containing polymetamorphic garnet. The hike thencontinues east for about 600 m through a side arroyo andinto Sand Canyon where the southern contact of the Priest
pluton is exposed. Stops in Sand Canyon will be made to
examine sericitized porphyroblasts of relict staurolite (stopb), granitic dikes crosscutting the pluton and its host rocks(stop c), and the Priest pluton itself (stop d).
General geology of the central and southern
Manzano Mountains
The central and southern Manzano Mountains consist of
two tectonometamorphic sequences separated by a ductile
shear zone. Greenschist-grade rocks are found to the northwest of the shear zone, and amphibolite-grade rocks arefound to its southeast (Fig. 1). Thrusting along the shear
zone emplaced a southern sequence (the amphibolite-grade
rocks) to the northwest over a northern sequence (thegreenschist-grade rocks) when P—T conditions were in thegreenschist facies. The southern sequence has been retrograded to the greenschist grade in the area around and southof the Priest pluton, probably synchronous with the thrusting. The following text summarizes the general geology ofthe northern, greenschist-grade sequence, the southern, amphibolite-grade sequence, and the portion of the southern
sequence that has been retrograded to the greenschist-grade.
It also discusses the Priest pluton and the ages of the amphibolite-grade and greenschist-grade metamorphic events.
Greenschist-grade sequence
Mineral assemblages in the northern tectonic sequence
confirm its assignment to the greenschist grade (Table 1). The
most common assemblage is quartz—muscovite—chlorite ±biotite ± tourmaline ± carbonate found in pelitic phyl-

74
lite. Aluminous phyllites contain quartz—muscovite—
chlorite ± kyanite ± chloritoid. A distinctive feldspathic
phyllite has alternating pink and green layers defined byvarying proportions of quartz, biotite, albite, muscovite, Kfeldspar, and chlorite. Rare mafic rocks contain chlorite—
albite—epidote—quartz ± actinolite. Quartzite contains
minor muscovite and hematite. Rare samples found near thecontact between the northern and southern sequences also
contain tiny (10-50 µm), nearly unzoned crystals of
spessartitic garnet.
Metamorphic P—T conditions for the northern tectonic
sequence were determined from mineral assemblages andgarnet—biotite geothermometry using the calibration of Williams and Grambling (1990). Mineral compositions wereobtained using the JOEL 733 electron microprobe in the
Department of Geology, University of New Mexico. Assemblages in mafic rocks require temperatures below about
500°C (cf. Apted and Liou, 1983; Moody et al., 1983). The
assemblage chloritoid ± kyanite indicates temperatures ofroughly 350-500°C, and the presence of kyanite indicatespressures above 2-3 kb. These constraints are consistent withgarnet—biotite geothermometry which indicates temperatures of 475-485°C.
Aluminous quartzite and phyllite collected near the contact between the northern and southern sequences containmylonitic foliation. Textures of kyanite, chloritoid, and garnet help constrain the timing of deformation and metamorphism. Kyanite grains are folded and kinked. Chloritoid
porphyroblasts are folded and/or aligned in the lineation
direction. Some large, randomly oriented, chloritoid porphyroblasts have grown across the mylonitic foliation. Individual mica grains in phyllite are terminated against
euhedral garnet grains. These mineral textures suggest thatthrusting was broadly synchronous with greenschist-grade
metamorphism.
All rock types in the greenschist-grade, northern tectonic
sequence (previously mapped as Blue Springs Schist andWhite Ridge Quartzite; cf. Stark, 1956; Bauer, 1982) areintensely foliated and lineated. Most pelitic phyllites exhibit
two or three foliations (for example, compositional layeringand two crosscutting grain-shape foliations). Other phyllites
contain boudinaged and folded quartz veins and quartz rods
(quartz veins stretched along foliation planes). Extension
lineations, which commonly appear on foliation surfaces, aredefined by aligned quartz or micas; crenulation lineations
also can be seen in most outcrops. Feldspathic quartzite hasalternating, centimeter-scale pink and green layers definedby varying amounts of biotite. Locally, these layers havebeen deformed into sheath folds or have been isoclinallyfolded, transposed, and refolded. Quartzite crops out as a 2040 m thick, massive white or pink layer with a foliation
defined by differently colored beds or by thin layers of
opaque minerals. Some isoclinal fold hinges can be seen
within foliation layers. Most quartzite is fine-grained, although some contains porphyroblasts of quartz deformed
into tectonic "fish," which generally indicate a top-to-thenorth sense of shear, in a finer-grained matrix. The quartzite
seems to have behaved more coherently than the phyllites.
Rocks are more intensely deformed closer to the contact
with the southern tectonic sequence. A distinctive feldspathic
mylonite crops out near the contact. Its strong, typicallysouth-dipping, mylonitic foliation is defined by alternating
black and pink, tan, or orange layers (<1 cm thick). Thisfoliation is folded in centimeter-scale folds, and a strong

extension lineation defined by aligned minerals can be seen
on most foliation surfaces. On joint surfaces parallel to thelineation and perpendicular to the foliation, quartz and plagioclase porphyroclasts have asymmetric "tails." The asymmetric structures in the mylonite generally indicate a toptothe-north sense of shear.
Amphibolite-grade sequence
Mineral assemblages observed in rocks in the southern
tectonic sequence are listed in Table 1. Metarhyolite contains
thin, semipelitic beds with garnet—biotite—quartz—plagioclase—K-feldspar. Mafic rocks typically have the assemblage
blue-green hornblende—plagioclase—chlorite—quartz ± epidote ± FeTi oxides. Pelitic schist contains garnet—biotitechlorite—muscovite—plagioclase—quartz ± ilmenite. Aluminous schists that crop out as thin layers within quarzite havethe assemblage quartz—muscovite—chloritoid—stauroliteAl2SiO5. The common form of the Al 2 SiO 5 in these schists iskyanite, but two specimens contain kyanite—andalusite
sillimanite. In these two rocks, kyanite forms bent, corroded
relicts inside andalusite and muscovite. Sillimanite occurs as
fibrolitic mats, with fibers aligned in the dominant foliation or
bent and folded, locally defining two foliations. Andalusite
forms large, undeformed porphyroblasts that have overgrown
the foliations defined by kyanite and sillimanite. Texturesindicate that andalusite was the last Al 2 SiO 5 poly-morph toform, but do not constrain the relative timing of kyanite andsillimanite growth.
Aluminum silicate triple-point assemblages suggest thatpeak P—T conditions of the southern sequence approached
4 kb, 500°C (cf. Grambling et al., 1989). Garnet—biotite
geothermometry (Williams and Grambling, 1990) and garnet—biotite—plagioclase geobarometry (Ghent and Stout,1981) are consistent with this P—T estimate. They indicatethat peak metamorphic conditions for the southern tectonic
sequence were near 540°C, 4 kb.
Deformation is pronounced in rocks of the amphibolitegrade tectonic sequence. Foliations generally trend
northeast and dip south, with extension lineations aligned
down-dip on foliation planes. Metarhyolite near thetectonic contact with the greenschist-grade sequence isfiner-grained and has a stronger foliation than elsewhere,suggesting that it is mylonitized. Foliation in metarhyolite
is crosscut by shear bands oriented about 20° from thedominant foliation. Tectonic "fish" of plagioclase and"snails" of plagioclase and quartz occur along foliation
planes. Asymmetric porphyroclasts and shear bands inmetarhyolite indicate a top-tothe-north sense of shear, thesame as that observed in adjacent greenschist-grade rocks.
Farther south, the metarhyolite becomes coarser-grained
so that individual quartz and feldspar grains are visible inhand specimen. The metarhyolite is schistose in places andcoarser-grained micas can be seen on foliation planes. Locally, it contains rare rounded cobbles of granitic rock (alsowith asymmetric tails) from approximately 2 x 2 cm up to 6
x 9 cm in size.
Interspersed within the metarhyolite are amphibolite podsor layers from less than I m to 60 m thick. Generally, themafic layers become thicker and more abundant farther southof the shear zone. Most thin amphibolite lenses can be traced
for no more than a few meters, whereas larger outcrops canbe traced visually across valleys and up adjacent ridges. In
larger outcrops, the amphibolite consists of very fine- andcoarse-grained layers separated by sharp contacts. Amphi-

75
bolite is foliated and, in some coarse-grained, foliated outcrops, clasts of finer-grained epidote, amphibole,plagioclase, and quartz have asymmetric tails of similarmaterial drawn out along foliation planes. Once again, theasymmetric tails indicate a top-to-the-north sense of shear.
Intergrowths of acicular hornblende and biotite (0.5 x 2 cm)
appear randomly oriented within the foliation planes ofsome amphibolites.
Pelitic schists have alternating quartz-rich and mica-richlayers, with the dominant foliation defined by aligned mica
grains. Quartz-rich layers contain micas that preserve an
earlier foliation approximately perpendicular to the dominant foliation. Abundant inclusions of quartz in garnet areparallel to the dominant foliation or have slightly sigmoidal
traces. This may indicate that the garnets are slightly rotatedor that they grew over a slightly folded foliation. Micaceous
layers are crenulated and crosscut by shear bands. The shear
bands indicate a top-to-the-north sense of shear.

the western margin of the Priest pluton. All lithologies are
fine-grained. Supracrustal rocks are crosscut by at least two
foliations, one that is defined by aligned micas subparallel
to compositional layering, and a second that is typically acrenulation cleavage. Phyllites generally have both an extension and crenulation lineation on foliation planes, theformer defined by aligned grains of muscovite and chlorite.
The chlorite phyllite exposed at the southern contact with
the Priest pluton consists of alternating phyllitic and quartzrich layers. The quartz-rich layers are discontinuous andrange from less than one centimeter to several centimetersin thickness. The compositional layers are disharmonicallyfolded. West of the southern margin of the pluton, supracrustal rocks are folded into a megascopic, upright, northplunging antiform. Most rocks exposed in this area form
the eastern limb of this fold; much of the western limb has
been removed by rift-related normal faults or covered byrift-related sediments.

Retrograded greenschist-grade sequence

The Priest pluton
The Priest pluton is composed of a main body of quartzmonzonite and lesser amounts of other granitic intrusive
rocks and dikes. The pluton typically contains coarsegrained quartz, plagioclase, muscovite, biotite, and epidote,
together with K-feldspar megacrysts and trace amounts oftitanite, apatite, and zircon. Near its southern margin, thepluton appears to be composed of several different granitic
rock types which vary in feldspar content and grain size. A
white, medium- to coarse-grained dike contains plagioclase
and no K-feldspar; it appears to crosscut the main body ofthe pluton. An equigranular, coarse-grained granitic portionof the pluton contains both plagioclase and K-feldspar.
There appears to be a compositional and textural gradation
between this portion of the pluton and the main body that
contains K-feldspar megacrysts. Also near the southern
margin, pink aplitic dikes crosscut the pluton and country
rocks. These dikes may represent late, separate intrusions or
the last stages of melt injected prior to complete
crystallization of the pluton.
The contact between the Priest pluton and the countryrock is intrusive, in places modified slightly by tectonism.
The interior of the pluton appears to be undeformed. The
southwestern and northernmost margins appear to be simple
intrusive contacts; the pluton contains xenoliths of countryrock, dikes of aplitic granite intrude the country rock, andthe igneous rocks are not foliated. The northwestern marginappears to be tectonized. The pluton near this margin is
weakly foliated with a subhorizontal north-trending extension lineation. Granitic dikes and sills in the country rockare strongly foliated. However, the appearance of xenolithsof adjacent country rock in this area, and of dikes and sillsof plutonic rock in nearby country rock, suggests that
motion along this contact was minimal. The eastern marginof the Priest pluton is faulted against Paleozoic sediments.

Retrograded greenschist-grade rocks are exposed in thesouthernmost Manzano Mountains in the Abo Pass area andinclude quartzite, pelitic phyllite, and quartzofeldspathic rock.
Mineral assemblages are reported in Table 1. Phyllites containthe assemblage quartz—muscovite—plagioclase—chlorite ±biotite ± garnet. Quartzofeldspathic rocks contain quartzplagioclase—K-feldspar—muscovite—biotite—garnet.
Garnet crystals in pelitic phyllites and quartzofeldspathic rocks
are very small (<1 mm) and have cloudy, irregular cores surrounded by clear rims 10-50 µm thick. Chlorite in chlorite
phyllite occurs as large, polycrystalline "porphyroblasts" that
contain numerous FeTi-oxide inclusions. Because biotite
accepts considerable titanium in its structure and chlorite does
not, the chlorite—FeTi oxide "porphyroblasts" are interpreted
to be pseudomorphs after biotite. Elsewhere, the phyllite
contains large, relict porphyroblasts that have been partly tocompletely replaced by sericite. Microscopic observations
show that locally sericite has replaced staurolite. This is shownby various amounts of replacement. One staurolite
porphyroblast (2 x 1 mm) is surrounded by a halo of sericite
about 0.2 mm thick. More commonly, relict staurolite occurs
as tiny remnants (0.2x 0.2 mm) inside a mat of sericite 0.5 x 1
cm. The mineral reaction textures described here aresuggestive of retrograde metamorphism accompanied by an
influx of hydrous fluid.
Minerals from a pelitic phyllite with unaltered garnetbiotite—chlorite—muscovite—plagioclase—quartz
wereused to determine metamorphic conditions during
retrogression in this area. Garnets have cloudy cores, withXgrs of 0.150.16, surrounded by clear rims. And abruptchange in Xgrs to 0.04-0.05 marks the transition to the clear
rims. Garnet cores are interpreted to represent growth during
the earlier, amphibolite-grade metamorphism and rims areinterpreted to represent growth during the later, retrograde,greenschistgrade thermal overprint. Biotite grains are small
and unaltered in this sample, as are plagioclase andmuscovite; both apparently grew or re-equilibrated during theretrograde greenschist-grade overprint. P—T conditions
during the second metamorphic event, determined fromgarnet(rim)—biotite—plagioclase—muscovite
thermobarometry (Williams and Grambling, 1990; Ghent andStout, 1981) were 485°C, 2.75 kb.
Within the retrograded southern sequence, outcrop patterns trend north-northwest to south-southeast, parallel toGeochronology
Muscovite and hornblende concentrates from samples
collected within the amphibolite-grade sequence have been
dated by 40Ar/39Ar methods to constrain the timing of deformation and metamorphism. Techniques follow those described by Dallmeyer and Keppie (1987) and Dallmeyer and
Gil-Ibarguchi (1990). Results are presented in more detailelsewhere (Thompson et al., in preparation). Blocking temperatures for argon diffusion are taken to be about 500 ± 25°C

76
for hornblende (Harrison, 1981) and about 400°C for muscovite, the latter based upon preliminary experimental data of
Robbins (1972) in diffusive equations of Dodson (1973,1979). Hornblende and muscovite argon ages are particularly
useful in this study because their blocking temperatures
bracket temperatures maintained during thrusting within theshear zone in Monte Largo Canyon. Consequently, hornblende ages likely record the age of the amphibolite-grade
thermal event that preceded thrusting, whereas muscovite
ages likely record cooling following thrusting.
Hornblende from sample 6 displays an internally discordant release spectrum (Fig. 6a) which defines a total-gas ageof 1448.8 ± 3.0 Ma. The apparent K/Ca ratio is constant over
94% of the argon release spectrum (Fig. 6b), indicating thatthe bulk of the experimentally derived gas evolved from acompositionally uniform population of intracrystalline"sites." Isotope data from the part of the release spectrum
that corresponds to a constant apparent K/Ca ratio (experimental temperatures between 850 and 1050°C), define an
isotope correlation (Fig. 6c) corresponding to an age of
1438 ± 5 Ma. This is interpreted to date cooling of the amphibolite-grade sequence through about 500°C.
Muscovite release spectra are internally discordant (Fig.
7). Sample 7 is the least discordant with low-temperatureages that systematically increase to define an intermediateto high-temperature plateau of 1361 ± 3 Ma (Fig. 7a). The
muscovite concentrate from sample 4, slightly more discordant, gives a high-temperature plateau age of 1338 ± 3
Ma (Fig. 7b). The discordant muscovite spectra are interpreted to result from a variable thermal overprint at about
1000-1100 Ma on intracrystalline systems that had initiallycooled through argon retention temperatures at around13401360 Ma. Because geothermometry and phaserelations suggest that the temperature of thrusting was
480°C (above the closure temperature for argon diffusion inmuscovite), we interpret the 1340-1360 Ma plateau ages torecord the timing of thrusting.

Conclusions
Proterozoic rocks in the central and southern Manzano
Mountains record two tectonometamorphic events recognized on the basis of metamorphic grade, structural relationships, and 40Ar/39Ar geochronology. 40Ar/39Ar ages
recorded by hornblende in the southern tectonic sequencesuggest that amphibolite-grade metamorphism occurred
around 1440 Ma. This is interpreted to represent the age ofthe first tectonometamorphic event, M 1 . Aluminum-silicate
triple-point assemblages and garnet—biotite—plagioclase-

77
muscovite thermobarometry indicate temperature and pressure conditions of 540°C, 4 kb for the peak of M, amphibolite-grade metamorphism.
40Ar/39Ar muscovite ages from the same tectonic
sequence were reset during a second tectonometamorphic
event, M2. The second event was synchronous with thrustingalong the Monte Largo shear zone that separates greenschistgrade rocks from amphibolite-grade rocks; argon isotopic
data from muscovite suggest that thrusting occurred around1350 Ma. Mineral assemblages and garnet—biotite
geothermometry of the northern sequence indicate
temperatures of 475-485°C and pressures of 2-3 kb for theupper greenschist-grade metamorphic
event,
M2.
Thermobarometry of rocks from the retrograded portion of
the amphibolite-grade sequence suggests that P—T
conditions during retrograde metamorphism were 480°C,
2.75 kb. The temperature for retrograde metamorphism is thesame as the temperature associated with M2 in Monte Largo
Canyon. The similarities in temperature between these twoareas suggest that M2 resulted in retrograde metamorphism of
the amphibolite-grade rocks and prograde metamorphism of
the greenschist-grade rocks. The retrogression is shown
clearly in the Abo Pass area, whereas the prograde
metamorphism is best seen north of Monte Largo Canyon.
Dallmeyer et al. (1990) summarized four tectonometamorphic events in Proterozoic rocks of New Mexico. Their
M 1 and M2 events are recorded over a large area in central
and northern New Mexico and correspond to M 1 and M2 inthe central Manzano Mountains. Dallmeyer et al. (1990)
reported that M 1 involved regional ductile extension accompanied by amphibolite-grade metamorphism. Their primaryevidence for M2 comes from the ductile thrust fault
discussed in this study. The age and character of M 1 and M2
are summarized in Table 2. M3 and M4 (Dallmeyer et al.,1990) are later thermal events dated at about 1300 Ma and9501100 Ma, respectively, most evident in parts of northern
New Mexico. They are represented in the present study areaby a partial thermal overprint on muscovite age spectra.

Albuquerque Country II: New Mexico Geological Society, Guidebook
33, pp. 211-216.
Dallmeyer, R. D., and Gil-Ibarguchi, J. I., 1990, Age of amphibolitic
metamorphism in the ophiolitic unit of the Morias allochthon (Portugal):
implications for Early Hercynian orogenesis in the Iberian Massif: Journalof the Geological Society of London, v. 147, pp. 873-878.
Dallmeyer, R. D., and Keppie, J. D., 1987, Polyphase late Paleozoic
tectonothermal evolution of the southwestern Meguma Terrane, Nova
Scotia: evidence from 40Ar/39Ar mineral ages: Canadian Journal of
Earth Sciences, v. 24, pp. 1242-1254.
Dallmeyer, R. D., Grambling, J. A., and Thompson, A. G., 1990, Age andcharacter of Proterozoic polymetamorphism in New Mexico (abs.):
Geological Society of America, Abstracts with Programs, v. 22, no. 7,
p. 113.
Dodson, M. H., 1973, Closure temperature in cooling geochronologicaland petrological systems: Contributions to Mineralogy and Petrology,
v. 40, pp. 259-274.
Dodson, M. H., 1979, Theory of cooling ages; in Jager, E., and Hunziker,
J. C. (eds.), Lectures in Isotope Geology: Springer-Verlag, Berlin, pp.
194-202.
Ghent, E. D., and Stout, M. V., 1981, Geobarometry and geothermometry ofplagioclase—biotite—garnet—muscovite assemblages: Contributions toMineralogy and Petrology, v. 76, pp. 92-97.
Grambling, J. A., Williams, M. L., Smith, R. F., and Mawer, C. K., 1989,
The role of crustal extension in the metamorphism of Proterozoic rocks
in New Mexico: Geological Society of America, Special Paper 235, pp.
87-110.
Harrison, T. M., 1981, Diffusion of 40Ar in hornblende: Contributions toMineralogy and Petrology, v. 78, pp. 324-331.
Kelley, V. C., 1982, Diverse geology of the Hubbell bench, Albuquerquebasin; in Grambling, J. A., and Wells, S. G. (eds.), Albuquerque Country
II: New Mexico Geological Society, Guidebook 33, pp. 159-160.
Kretz, R., 1983, Symbols for rock forming minerals: American Mineralogist, v. 68, pp. 277-279.
Kudo, A. M., 1982, Rift volcanics of the Albuquerque basin: overview
with some new data; in Grambling, J. A., and Wells, S. G. (eds.),
Albuquerque Country II: New Mexico Geological Society, Guidebook
33, pp. 285-290.
Machette, M. N., 1982, Quaternary and Pliocene faults in the La Jencia
and southern part of the Albuquerque and Belen basins, New Mexico:
evidence of fault history from fault-scarp morphology and Quaternary
geology; in Grambling, J. A., and Wells, S. G. (eds.), AlbuquerqueCountry II: New Mexico Geological Society, Guidebook 33, pp. 161170.
Moody, J. B., Meyer, D., and Jenkins, J. E., 1983, Experimental characterization of the greenschist/amphibolite boundary in mafic systems:
American Journal of Science, v. 283, pp. 48-96.
Stark, J. T., 1956, Geology of the southern Manzano Mountains, New
Mexico: New Mexico Bureau of Mines & Mineral Resources, Bulletin
34, 46 pp.
Thompson, A., Grambling, J. A., and Mawer, C. K., 1989, Proterozoic
tectonic assembly of southwestern North America: evidence from ductile
shear zones in the Manzano Mountains, New Mexico (abs.): GeologicalSociety of America, Abstracts with Programs, v. 21, no. 6, p. 215.
Thompson, A., Grambling, J. A., Mawer, C. K., Dallmeyer, R. D., andDaniel, C. G. (in preparation), Multiple tectonothermal events in Proterozoic rocks of central New Mexico.
Vernon, R. H., Mawer, C. K., and Grambling, J. A. (in preparation), The
1.45 Ga Sandia Granite, New Mexico: a syntectonic, high-level, Proterozoic batholith.
Williams, M. L., and Grambling, J. A., 1990, Manganese, ferric iron, andthe equilibrium between garnet and biotite: American Mineralogist, v.
75, pp. 886-908.

79

New Mexico Bureau of Mines & Mineral Resources Bulletin 137, April 1991

Late Pennsylvanian stratigraphy and paleontology of the Kinney Brick
Quarry, Manzanita Mountains, New Mexico
Spencer G. Lucas and Phillip Huber
New Mexico Museum of Natural History, P.O. Box 7010, Albuquerque, New Mexico 87194

Introduction
Pennsylvanian—Permian strata are widely exposed in
the Manzanita and Manzano Mountains of central New
Mexico (Fig. 1). These rocks pertain to the Sandia
Formation, Madera Group, and Abo Formation and range inage from Atokan to Wolfcampian (Fig. 2). The Kinney
Brick Quarry is a clay pit developed in the lower Virgilian
Pine Shadow Member of the Wild Cow Formation of theMadera Group, about 8 mi south of Tijeras (Fig. 1). Strata
exposed here provide a unique glimpse of a diverse andwell-preserved Late Pennsylvanian lagoonal biota. Here, wesummarize previous research on this biota, describe its
stratigraphic context, delineate fossil assemblages, andpresent a preliminary interpretation of the depositional
environments of the strata exposed at the quarry.
Previous researchStudents at the University of New Mexico discovered therich fossil biota of the Kinney Quarry in 1961 (B. S. Kues,
written comm. 1990). During the 1960's, large collections
were made of fishes, plants, and other fossils from thequarry, especially by D. Dunkle and S. Mamay of the National Museum of Natural History. Others who have madesignificant collections at the quarry include D. Berman, J. P.
Bradbury, R. Burton, M. D. Gottfried, P. Huber, K. Kietzke,
B. Kues, N. LaFon, T. Lehman, and F. Schram.

Extensive fossil collections from the Kinney Brick Quarry
are now housed at the Carnegie Museum of Natural History,
National Museum of Natural History, New Mexico Museumof Natural History (hereafter NMMNH), University of Kansas Museum of Natural History, and University of New
Mexico.
Stratigraphic data on the quarry have been presented by80
Stukey (1967), Gottfried (1987b), and Huber et al. (1989).
Kelley and Northrop (1975) provided a list of the fossil biotaof the quarry, although most of the taxa they listed have yet tobe adequately documented. Published descriptions oforganisms are: megafossil plants (Mamay, 1981, 1990; Ashand Tidwell, 1982), conodonts (Kelley and Northrop, 1975),
insects (Carpenter, 1970), malacostracans (Schram andSchram, 1979; Schram, 1981), the pectinacean bivalve
Dunbarella (Clark, 1978), eurypterids (Kues, 1985), fishes(Zidek, 1975; Gottfried, 1987a, b), and amphibians(Berman, 1973; Hunt et al., 1989). A variety of other fossils
known from the quarry that have not been described are also
listed here (Table 1), but no comprehensive review of theKinney fossil biota has yet been published.
Stratigraphy and age
The quarrying operation has exposed about 28 m of theupper part of the Pine Shadow Member of the Wild Cow
Formation (Fig. 3). In the Manzanita and Manzano Mountains, the Pine Shadow Member consists of 30-50 m of

intercalated limestone, calcareous shale, and sandstone(Myers, 1973, 1982). Stukey (1967) estimated that the baseof the Wolfcampian Abo Formation is stratigraphically about
100 m above the Kinney Quarry exposures.
The active quarry floor typically is developed on black,platy and laminar micrite with a high content of black clay
(Fig. 4, unit 1). The upper 10 cm of this limestone arefinely laminated and grade into the overlying shale (Fig.
4, unit 2).
The basal shale exposed at the quarry is 0.4 m thick,olive gray to olive black, well indurated, calcareous andhighly fossiliferous (Fig. 4, units 3-4). Part of this interval
is highly fissile and not calcareous. Above it is 2.0 m ofolive-gray, calcareous shale that contains thin bands ofplastic clay (Fig. 4, units 5-7). Individual laminae in theseshales are continuous on strike for at least 10 m.
Above, and in gradational contact with the underlying
shale, are 10.7 m of olive-gray and greenish-gray, laminarcalcareous silty shale that are moderately bioturbated (Fig.
4, units 8-9). These strata coarsen upward slightly and aremarked by a rapid decrease in floral and faunal diversity
near their base. Above them are 12.7 m of olive-gray, silty
shale and claystone interbedded with lenticular, yellowishorange and grayish-brown, subarkosic/sublitharenitic, laminar and ripple-laminar sandstone ledges (Fig. 4, units 1016). The section is completed by 1.0+ m of limestonecobble conglomerate with clasts as much as 8 cm indiameterandgrayish-yellow,
trough-crossbedded
subarkosic—sublithic arenite (Fig. 4, unit 17).
Some early workers (e.g. Stukey, 1967; Berman, 1973)

assigned a Permian age to the Kinney fossils, based primarily on an unsubstantiated report of the Permian indexfossil Callipteris (cf. Read and Mamay, 1964) from thequarry (Kelley and Northrop, 1975). However, fusulinids of
the Pine Shadow Member, from more open-marine facies
than are those exposed at the quarry, include Triticites species indicative of the early Virgilian (Myers, 1982, 1988).
Therefore, the Late Pennsylvanian age of the Kinney fossils
seems to be well established.
Fossil assemblages
Almost all published fossils from the Kinney Quarry werecollected from units 2-4 of our measured section (Fig. 4). We
have sampled these units extensively and have obtainedmuch smaller collections from other units at the quarry. All
fossils we collected are housed at the NMMNH. In addition,
B. S. Kues has extensively sampled units 1, 5, and 15 of ourmeasured section; his collections are housed at the Universityof New Mexico. Based on Kues' and our collections, werecognize two floral and four faunal assemblages at the
Kinney Quarry (Fig. 4).
1. Pteridosperm floral assemblage: The flora from units
1-4 consists of calamitaleans, pteridosperms, cordaitaleans,
and Walchia. Calamites occurs in the basal micrite (unit 1)with Asterophyllites and Annularia, followed in abundanceby pteridosperms and cordaitalean leaves. Pteridosperms
are very common in the overlying shale (units 3-4) and are
typified by Neuropteris (Fig. 5A). Calamitaleans still arepresent, as are leaves of Cordaites. One or more species ofWalchia are present, though not abundant.

84
2. Walchia floral assemblage: This assemblage consists
almost entirely of walchian conifers (Fig. 5B), whose leafy
shoots are abundant throughout units 8 and 9. The only
other plant we recovered from unit 9 is a single, large
branch of Pecopteris.
3. Bivalve—Lingula
faunalassemblage:This
assemblage is confined to the basal micrite (unit 1) and isdominated by the bivalves Solemya and Myalina, and by theinarticulate brachiopod Lingula. Subsidiary members of thisassemblage
includeunornamented
ostracodes,
conchostracans, the articulate brachiopod Chonetinella,Dunbarella, and small ammonoids. Other bivalves,brachiopods, gastropods, and orthocerid nautiloids areuncommon, and bryozoan and crinoid fragments are rare.
Conodonts, insect body parts, and disarticulated fish
remains also are rare (Kues, 1991, in press).
4. Fish faunal assemblage: The uppermost micrite andoverlying 0.4 m of shale (units 2-4) contain a fish-dominated assemblage which includes elasmobrachs, dipnoans,coelacanths, acanthodians, and actinopterygians (Fig. 5E).
In terms of diversity and number of specimens, actinopterygians dominate the fish assemblage. This is the main collecting level at the quarry from which almost all of thepreviously reported fauna has been collected. The diverse
fauna also includes spirorbid and other worms, arthropods(conchostracans, ostracodes, malacostracans [Fig. 5F], myriapods, insects [Fig. 5D]), small Dunbarella, gastropods
and allochthonous nautiloids, brachiopods and fenestratebryozoans, as well as rare trimerorhachid amphibians.
Spirorbid worms, ostracodes, insects, small Dunbarella,gastropods, and fishes occur throughout units 3 and 4. Spirorbid worms are represented by spiral tubes often attached
to floral elements of the pteridosperm assemblage. The ostracodes and gastropods are generally poorly preserved,
crushed, and lacking significant morphological detail. Small
Dunbarella are common, preserved as compressed internalmolds or recrystallized, disarticulated valves. The clearly
allochthonous elements of this assemblage are crushed, disarticulated valves of articulate brachiopods, incomplete orthocerid shells, and small fragments of fenestrate bryozoans.
Worms, millipedes, insects, and fish usually occur as carbon
films. Articulated acanthodians and actinopterygians arecommon, whereas the dipnoans and coelacanths are mostly
represented by isolated scales and dermal bones. Complete
skeletons of xenacanthid elasmobranchs are rare and restricted to the lower part of unit 3, within 5 cm of the contact
with the underlying shaly limestone (unit 2). Conchostracans
are very abundant in the lower 8-10 cm of unit 3 and occur
as isolated individuals or clusters of 10 or more carapaces.
Shrimp are restricted to a narrow, 4 cm thick zone just above
the conchostracan-bearing horizon.
5. Large Dunbarella-dominated assemblage: The transition to this assemblage, which is present in units 5-7,
occurs over a vertical distance of 8 cm. Unusually large
specimens of Dunbarella (Fig. 5C) dominate this assemblage, which also includes a few fishes. Rare eurypterids
occur only here. Articulated and disarticulated, recrystallized valves of Dunbarella in all stages of growth abound,and individuals are restricted to specific laminae, possibly
due to seasonal fluctuations (Clark, 1978).
6. Small Dunbarella assemblage: This consists of small
valves of Dunbarella and, in unit 15, less common Lingulaand Myalina. It is associated with the Walchia floral assemblage sporadically in the silty shales of units 8-9. Above

unit 9, shale units generally are unfossiliferous except forlarge pieces of carbonized wood and badly macerated conifer debris. However, unit 15 contains the small Dunbarella
assemblage and moderate amounts of plant material.
Depositional environments
Several authors have suggested depositional environments for the highly fossiliferous, fine-grained rocks low in
the quarry section (Fig. 4, units 1-7), including nearshore
marine (Gottfried, 1987a, b), brackish marine—lagoonal(Berman, 1973; Zidek, 1975; Schram, 1981), and brackishlagoonal with a significant fresh-water influence (Clark,
1978; Kues, 1985). All of these interpretations arereasonable, depending on which floral/faunal assemblages(and, consequently, which aspect of the Kinney strata) arestudied. Combining our physical stratigraphy with the fossil
assemblages, we recognize several distinct depositional
environments at the Kinney Quarry (Fig. 4). These make up
a regressive sequence of the "R2" type described by Smith(1989), in which limestone grades up through prodelta/deltafront clastics with a capping delta-plain facies.
The basal micrite (units 1-2) represents deposition in anearshore marine environment that received some freshwater/clastic sediment input. We point to its lithology (noteespecially the high black-clay content of this micrite) andfauna (note articulate brachiopods such as Chonetinella,characteristic of more open-marine-shelf facies of the Pine
Shadow Member) to support this conclusion. The blackclay content, terrestrial plant debris, and euryhaline elements of the fauna (especially abundant Lingula) are consistent with deposition near the shoreline with a significant
fresh-water input.
The overlying highly fossiliferous shales (units 3-4) were
deposited in a calm lagoonal or estuarine environment witha significant fresh-water input. Uniform, fine grain size,
fine lamination and lack of bioturbation, dark colors, andpreservation of soft-bodied forms suggest deposition inquiet, oxygen-poor waters with restricted circulation. A
lowland, pteridosperm-dominated flora and fresh-water
faunalelements(especiallyconchostracansandtrimerorhachid amphibians) suggest low salinity.
Overlying shales (units 5-7) represent a similar facies,
but probably with a greater fresh-water influence. Dominantelements are Dunbarella, which must be euryhaline (cf.
Clark, 1978), and terrestrial plants.
Overlying silty shales (units 8-9) are interpreted by us asrepresenting increased sedimentation rates in an estuarine
environment, brought about by the onset of a significant
fluvial discharge. The Walchia floral assemblage abounds
here, and there are sparse, small Dunbarella. Most of thefloral assemblage, typically interpreted as an inland, relatively xerophytic flora (Read, 1947; Ash and Tidwell,1982), may have been floated into an environment
characterized by frequent shallow ponding and deposition
on a better drained surface than the underlying shales.
Overlying laminar/ripple-laminar sandstone ledges andintercalated shales/claystones (units 10-14) are interpretedby us as delta-front, distributary mouth bars, and associateddeposits. Unit 15 is a shale that shows marine influence
indicated by the presence of Lingula and Myalina. This unit,
and the overlying fluvial sandstone/conglomerate, may be
the base of another transgressive sequence.
In conclusion, we interpret the stratigraphic sequence at
the Kinney Quarry as mostly reflecting a marine regression

85
corresponding to the progradation of a clastic delta. Shifting
sediments from the delta probably isolated an embayment
from normal marine conditions as a clastic wedge developedand extended seaward. Lagoonal conditions were establishedand clastic input was initially restricted to clay-size particles.
Eventually, the embayment was filled by silty shales from anadvancing delta plain that later deposited sand. The onset ofa subsequent transgression is documented by the highest
strata in the quarry section.
Access to the Kinney Brick Quarry
To reach the Kinney Brick Quarry from Albuquerque,
proceed east 13.9 mi on 1-40 from the 1-40/1-25
interchange to Exit 175, Tijeras-Cedar Crest. From the exit
continue south on NM-337 (formerly NM-14) 7.6 mi (fromthe Tijeras Post Office) to a relatively inconspicuous dirt
road on the right marked by a mailbox just south of the U.S.
Forest Service Deadman Picnic Ground. This dirt road isthe entrance to the Kinney Brick Quarry, which is privately
owned and thus should not be entered without permission.
The quarry is located in the SE1/4, sec. 18, T9N, R6E,
Bernalillo County. Hawley et al. (1982, pp. 8-10) provided
a road log of the route from Tijeras to the quarry. The
highway mostly traverses outcrops of the Los Moyos
Formation of the Madera Group (Fig. 1).
Acknowledgments
We thank B. Jurgena and the Kinney Brick Company
for access to the quarry, M. Gottfried, S. Hayden, B. Kues,
J. Lorenz, G. Smith, K. Toolson, and J. Zidek for diverse
help, and B. Kues, J. Lorenz, and J. Zidek for their reviewsof this paper.
References
Ash, S. A., and Tidwell, W. D., 1982, Notes on the upper Paleozoic plantsof central New Mexico: New Mexico Geological Society, Guidebook
33, pp. 245-248.
Berman, D. S., 1973, A trimerorhachid amphibian from the Upper Pennsylvanian of New Mexico: Journal of Paleontology, v. 47, pp. 932-945.
Carpenter, F. M., 1970, Fossil insects from New Mexico: Psyche, v. 77,
pp. 400-412.
Clark, G. R., II, 1979, Byssate scallops in a Late Pennsylvanian lagoon(abs.): Geological Society of America, Abstracts with Programs, v. 10,
p. 380.
Gottfried, M. D., 1987a, A Pennsylvanian aeduelliform (Osteichthyes,
Actinopterygii) from North America with comments on aeduelliform
relationships: Palaontologische Zeitschrift, v. 61, pp. 141-148.
Gottfried, M. D., 1987b, A new long-snouted actinopterygian fish fromthe Pennsylvanian of north-central New Mexico: New Mexico Journalof Science, v. 27, pp. 7-19.

Hawley, J. W., Foster, R. W., Broadhead, R., and Love, D. W., 1982,
Road-log segment I-B: Tijeras Canyon to Abo Canyon via Estancia andManzano: New Mexico Geological Society, Guidebook 33, pp. 8-30.
Huber, P., Lucas, S. G., and Hayden, S. N., 1989, Stratigraphy, depositional environments and fossil assemblages of an Upper Pennsylvanianlagoonal deposit, central New Mexico: New Mexico Journal of Science,v. 29, p. 117.
Hunt, A. P., Lucas, S. G., and Huber, P., 1989, Trimerorhachid amphibianfrom the Late Pennsylvanian of central New Mexico (abs.): Journal of
Vertebrate Paleontology, v. 9, supplement, p. 26A.
Kelley, V. C., and Northrop, S. A., 1975, Geology of Sandia Mountains
and vicinity, New Mexico: New Mexico Bureau of Mines & Mineral
Resources, Memoir 29, 136 pp.
Kues, B. S., 1985, Eurypterids from the Wild Cow Formation (Upper
Pennsylvanian), Manzano Mountains, New Mexico: New Mexico Journal of Science, v. 25, pp. 23-31.
Kues, B. S., 1991, A Late Pennsylvanian Lingula-Solemya-Myalina community, Kinney quarry, Manzanita Mountains, New Mexico (abs.): New
Mexico Geology, v. 13, no. 1.
Kues, B. S. (in press), A Late Pennsylvanian (Virgilian) restricted marine
fauna from the Kinney Quarry, Manzanita Mountains, New Mexico:
New Mexico Bureau of Mines & Mineral Resources, Bulletin 138.
Mamay, S. H., 1981, An unusual new species of Dicranophyllum
Grand'eury from the Virgilian (Upper Pennsylvanian) of New Mexico,
U.S.A.: The Palaeobotanist, v. 28-29, pp. 86-92.
Mamay, S. H., 1990, Charliea manzanitana, n. gen., n. sp., and otherenigmatic parallel-veined foliar forms from the Upper Pennsylvanian of
New Mexico and Texas: American Journal of Botany, v. 77, pp. 858866.
Myers, D. A., 1973, The upper Paleozoic Madera Group in the Manzano
Mountains, New Mexico: U.S. Geological Survey, Bulletin 1372-F, 13
pp.
Myers, D. A., 1982, Stratigraphic summary of Pennsylvanian and Lower
Permian rocks, Manzano Mountains, New Mexico: New Mexico Geological Society, Guidebook 33, pp. 233-237.
Myers, D. A., 1988, Stratigraphic distribution of some fusulinids from theWild Cow and Bursum Formations, Manzano Mountains, New Mexico:
U.S. Geological Survey, Professional Pape' 1446-B, 39 pp.
Read, C. B., 1947, Pennsylvanian floral zones and floral provinces:
Journal of Geology, v. 55, pp. 271-279.
Read, C. B., and Mamay, S. H., 1964, Upper Paleozoic floral zones andfloral provinces of the United States: U.S. Geological Survey, Professional Paper 454-K, 35 pp.
Schram, F. R., 1981, Late Paleozoic crustacean communities: Journal of
Paleontology, v. 55, pp. 126-137.
Schram, F. R., and Schram, J. M., 1979, Some shrimp of the Madera
Formation (Pennsylvanian), Manzanita Mountains, New Mexico: Journal of Paleontology, v. 53, pp. 169-174.
Smith, G. A., 1989, Eustatic signals in cyclic deltaic to shallow-marine
clastics and carbonates, Sandia Mountains, New Mexico (abs.): Geological Society of America, Abstracts with Programs, v. 21, p. 145.
Stukey, A. H., Jr., 1967, Stratigraphic relations of the PennsylvanianPermian strata, Manzanita Mountains, New Mexico: Unpublished MS
thesis, University of New Mexico, Albuquerque, 64 pp.
Zidek, J., 1975, Some fishes of the Wild Cow Formation (Pennsylvanian)
Manzanita Mountains, New Mexico: New Mexico Bureau of Mines &Mineral Resources, Circular 135, 22 pp.

See next page for Appendix

Comment
Although cockroaches do occur in the Kinney deposits, the insect illustrated in Fig. 5D (p. 83) is not one of them. According
to J. Kukalova-Peck, Carleton University, Ottawa (written comm. March 1991), the specimen most likely is a megasecopteranshowing the thorax and head with a sucking rostrum. (Eds.)

New Mexico Bureau of Mines & Mineral Resources Bulletin 137, April 1991

87

A tale of two volcaniclastic aprons: Field guide to the sedimentology
and physical volcanology of the Oligocene Espinaso Formation and
Miocene Peralta Tuff, north-central New Mexico
Gary A. Smith', Daniel Larsen', Steve S. Harlan', William C. McIntosh', Daniel W. Erskine',and Steve Taylor''Department of Geology, University of New Mexico, Albuquerque, New Mexico 87131: 'New Mexico Bureau of Mines & Mineral Resources. Socorro, New Mexico 87801

Geologic background to excursions to the Espinaso Formation and Peralta TuffGary A. Smith
Introduction
Volcanism has played an important role in the Cenozoic
history of New Mexico. Volcanic fields of variable size,
eruptive character, and magma composition developed overlarge areas of the state. Volcaniclastic aprons, composed ofprimary, volcanic fragmental deposits as well as secondary,reworked sedimentary facies accumulated adjacent to ventareas and peripheral to the volcanic fields. The purpose ofthis excursion (Fig. 1) is to compare and contrast the
primary and secondary volcaniclastic strata associated withvolcanic centers of different compositions, tectonic settings,
and ages. Emphasis is placed on relatively proximal aprons,
which include alluvial and pyroclastic-flow fans. In addition
to descriptions of the field-trip stops, three short papers
illustrate the utility of integrating paleomagnetic, petrologic,
and geochronologic methods with more obvious
sedimentologic and volcanologic studies in the investigationof volcaniclastic sequences.
Espinaso Formation
The Oligocene Espinaso Formation comprises the erosional remnants of widespread aprons of coarse-grained volcaniclastic debris that accumulated adjacent to a long-active
chain of volcanic-vent complexes stretching 40 km southward from near the southwestern edge of Santa Fe (Fig. I).
The vents are marked by stocks, laccoliths, sills, and dikesof the Ortiz porphyry belt (OPB; Fig. 1). Level of structural
exposure rises northward so that preservation of volcanic
rocks and volcaniclastic debris, assigned to the Espinaso
Formation, is restricted to the northern end of the belt. Twoprincipal areas of Espinaso Formation exposure occur, thetype section at Espinaso Ridge and a larger area near thetowns of Cerrillos and La Cienega (Fig. 2A). Espinaso Ridge
detritus was derived from the Ortiz Mountain eruptive
center, whereas the section east of Cerrillos is related tovolcanism in the Cerrillos Hills.
Two petrographic suites are associated with OPB magmatism (Maynard et al., 1990). Early calc-alkaline latites
and quartz latites were erupted between about 36 Ma and
34 Ma and were followed by alkaline latites, and consanguineous gold-mineralized monzonites, between about 30
and 28 Ma (Baldridge et al., 1980; Kautz et al., 1981;
Maynard et al., 1990). Although few lava flows or primary
pyroclastic deposits remain, the composition of Espinaso
Formation conglomerate clasts and sandstones faithfullyrecords the two episodes of volcanism (see Smith and Erskine, this volume). The products of the two eruptive suites
are easily distinguished in the field by the dominance ofhornblende or hornblende + clinopyroxene as the ferromagnesian silicates in the calc-alkaline lithologies; whereas

(Bachman, 1975). At Espinaso Ridge, a calcic paleosol caps
the Espinaso Formation and is disconformably overlain byearly Miocene Santa Fe Group. There is no conclusive evidence for continued subsidence of Laramide basins or initiation of extensional rift basins during Espinaso deposition.
The volcaniclastics accumulated as aprons of debris surrounding the volcanic centers that choked adjacentdrainages with detritus during a period of relative structural
stability, during the transition from compressional toextensional deformation.
The Espinaso Formation consists of proximal, debrisflow dominated alluvial-fan deposits, best seen east of
Cerrillos, that grade distally into braidplain facies, which
are well exposed on Espinaso Ridge. Lava flows are rare,
but do occur at both localities. Pyroclastic-fall deposits
also occur sparsely in both sections and three ignimbrites
are exposed at Espinaso Ridge. Espinaso Formation latite
clasts are remarkably porphyritic (45-50% phenocrysts),suggesting viscous extrusions in the form of domes, whichwere then the sources for numerous block-and-ash flows
and avalanches of carapace debris that supplied abundant
lithicpyroclastic and autoclastic fragments to adjacentalluvial fans and braidplains.
The Peralta Tuff Member of the Bearhead Rhyolite

hornblende is nearly absent in the alkaline rocks, whereaugite occurs alone or with biotite.
OPB magmatism occurred during the tectonic transition
between the Laramide orogeny and mid-Tertiary extension.
In most locations, the Espinaso Formation rests conformably
upon, and grades into, the Galisteo Formation, a largely
Eocene clastic unit derived from Laramide orogenic highlands. This relationship indicates continuity of deposition in
Laramide basins into the initial stages of mid-Tertiary
volcanism. At Arroyo Hondo, south of Santa Fe, alkaline
Espinaso Formation debris, capped by a 29 Ma latite flow
(Baldridge et al., 1980), records east-flowing streams thatcrossed an erosion surface cut on Precambrian rocks of theLaramide Sangre de Cristo uplift (Fig. 1). The Espinaso
Formation at Cerrillos is overlain, along a marked angular
unconformity, by Pliocene(?) Santa Fe Group sediments

Although best known for great episodes of Quaternary
rhyolitic-ignimbrite volcanism and caldera formation, theJemez Mountains contain a large volume of Miocene andPliocene volcanic rocks, of diverse composition, and locallywell preserved volcaniclastic aprons (Gardner et al., 1986;
Waresback and Turbeville, 1990). During a 1.0 to 1.5 m.y.
period in the late Miocene, high-silica rhyolite was erupted
from at least a score of vents dispersed over the southern
and, to a lesser extent, northern parts of the Jemez Mountains. The resulting plugs, domes, and lava flows mark thevents for the Bearhead Rhyolite. Eruptions of flows anddomes were commonly preceded by phreatomagmatic explosions that constructed tuff rings and adjacent fan-shaped
accumulations of pyroclastic-flow and fall deposits. Alongthe southeastern margin of the volcanic field, these primary
pyroclastic deposits are intercalated with sheetflood andbraided-stream-dominated alluvial-fan facies. The volcaniclastic deposits, which are at least 275 m thick, are collectively known as the Peralta Tuff Member of the Bearhead
Rhyolite (Fig. 2B). Subsidence of the Rio Grande rift, thewestern bounding structures of which were active in theJemez Mountains during Peralta Tuff deposition, partly accommodated the volcaniclastic fans along the southeastern
margin of the volcanic field (Fig. 1).
Compared to the Espinaso Formation, the Peralta Tuff
contains a much greater volume of primary pyroclastic deposits, whereas debris-flow facies are much less common.
The more sandy nature of Peralta Tuff sediments may, inpart, reflect the greater availability of relatively fine-grained
pyroclastic debris. The Peralta Tuff also contains eolian tuffs
resulting from wind reworking of ash-fall deposits (Smith
and Katzman, in press).
AccessMany of the stops described in this guide are on private
land or public land with restricted access. Stop 1 is locatedin a restricted Bureau of Land Management (BLM) habitat
management area reached by crossing private property. Permission for access must be obtained from the ranch head-

89
quarters mentioned in the directions to Stop 1 (Fig. 2A).
Stop 3 is partly located on private property, although representative exposures can be seen by walking through thesection in Canada de la Cueva (Fig. 5), on BLM land. Stop
5 is located on private land, currently owned by Mr. Fred

Rivera of Pena Blanca, New Mexico. Stop 6 is located onland owned by the University of New Mexico. Inquiry foraccess should be directed to the author or to Mr. Fred
Dixon, caretaker, at Rancho Canada, north of Cochiti Lake,
New Mexico.

Stratigraphy and sedimentology of Espinaso Ridge (Stop 1)
Gary A. Smith
Directions to Arroyo del Tuerto (Stop 1)
From either Albuquerque or Santa Fe, travel 1-25 to the
Budaghers Exit (exit 257) and turn east. A frontage roadheads southward along the east side of the freeway, becoming a dirt road in 0.5 mi as it enters private property. Continue 0.9 mi to a ranch headquarters where it is necessary toobtain permission to proceed. Cross the wash, turn west,
and follow the main track south along the west side ofEspinaso Ridge. The route from here is on unimproved dirt
roads, requiring modest vehicle clearance and four-wheel
drive whcn wet. At 3.0 mi from the ranch headquarters theroad enters BLM land near a windmill. Proceed another 2.0
mi to another gate, taking the right fork just past the gate.
Continue 0.7 mi to the ford across Arroyo del Tuerto. At
this point, four-wheel-drive vehicles may proceed approx

imately 0.8 mi up the wash to Stop 1. An alternative routecrosses the wash and continues over rough road for an additional 0.9 mi where a short hike to the northeast leads to
the Arroyo del Tuerto section (Fig. 3).
Stratigraphy of Espinaso Ridge
The exposures in Arroyo del Tuerto were designated thetype section of the Espinaso Formation (then Espinaso Volcanics) by Stearns (1953). The northwest-striking ridge isheld up by resistant, mostly gray volcaniclastic strata thatoverlie yellow and red lithic-arkosic clastics of the Eocene
Galisteo Formation and are overlain to the northeast by Santa
Fe Group sediments. The abrupt termination of the ridge near
1-25 marks the trace of the San Francisco fault zone, a Rio
Grande rift structure (Fig. 2A). The Espinaso For-

90
mation is approximately 400 m thick and includes a biotitelatite lava flow that forms the high points along thenorthwest end of the ridge.
Kautz et al. (1981) interpreted the Espinaso Ridge exposures as coalesced alluvial-fan deposits extending from theOrtiz Mountains and Cerrillos Hills eruptive centers. Several
observations argue against this interpretation. Facies
distribution and grain size are very continuous and consistentalong the entire 13 km length of the ridge (Kautz et al.,1981), which would not be expected if the outcrops included
deposits derived from two different sources, and are notcompatible with abrupt facies and grain-size variations typical of alluvial-fan strata. The distance between the source
regions and Espinaso Ridge, approximately 20 km, would
also require an unusually large fan. Facies types and distribution are more characteristic of a braidplain (Smith,1987). Mineralogy and geochemistry of clasts collected from
Espinaso Ridge are different from those found east of
Cerrillos, suggesting that nearly all, if not all, the detritus at
Espinaso Ridge originated at the Ortiz Mountains eruptive
center (Smith and Erskine, this volume).
Stratigraphy and sedimentology of the
Arroyo del Tuerto section
Fig. 3 illustrates and describes the highlights of the Espinaso Formation to be viewed by a traverse through thetype section in Arroyo del Tuerto. A prominent yellow,
pebbly sandstone containing petrified wood occurs high in
the Galisteo Formation and is a useful regional marker,
which will be seen again near Cerrillos. Volcanic grains first
appear above the yellow sandstone recording the initiation of
OPB volcanism while this area was still largely a braid-plain
draining south from Laramide highlands. Because of thegradation between Galisteo and Espinaso lithologies, there islittle agreement on where the contact should be placed(Stearns, 1953; Lucas, 1982); I define it at the base of a
distinctive light-green, highly tuffaceous mudstone that
contains abundant titanothere bone fragments. Above thislevel the section is dominantly volcaniclastic, although acrystalline provenance is recorded by trace constituents
through another 50 m.
The section is described in five intervals in Fig. 3. Alternating depositional regimes in interval 2 (Fig. 3C) may
record episodic syneruption and inter-eruption sedimentation
as described by Smith (1987; in press). Ignimbrites (Fig. 3B)
and coarse-grained debris-flow deposits dominate theculminating deposition of calc-alkaline detritus in interval 3.
This interval is capped by a monolithologic hornblende-latite
breccia with an ash matrix (Fig. 3E) that probably represents
a block- and ash-flow deposit; this interpretation is supported
by rock-magnetism studies (see Harlan, this volume). Clasts
in this breccia are petrologically notable for recording thecommingling of characteristic latite magma with a calcalkaline lamprophyre (spessartite) melt. The lamprophyreoccurs as individual black clasts and as black inclusions
within gray latite blocks. The inclusions typically exhibit
irregular cuspate, fine-grained chill margins indicating
incomplete mixing of partly crystallized melts rather thanincorporation of accessory xenoliths within the latite. An
irregular erosion surface, with as much as 15 m of relief,
forms a disconformity above the block- and ash-flow deposit.
Above this surface is 15-20 m of relatively well sorted, wellrounded conglomerate (interval 4) similar to that seen near
the base of the Espinaso Formation (interval 1). Clast
compositions are markedly polylithologic compared to thosein the underlying 100 m of section (interval 3) and include
the first augite latites and biotite latites of the alkaline suite.
The erosion surface and conglomerate probably represent
initial degradation followed by the reestablishment of agravelly braided stream during the hiatus between eruption of
the compositionally distinctive magmas (a poorly constrained
period, perhaps as great as 6 m.y. in duration) and during theinitial rejuvenation of eruptive activity. Interval 5 consists offiner-grained, poorly sorted, scour- and fill-bedded pebbly
sandstones indicating deposition by flashy discharge, poorly
confined, shallow braided streams choked with sandy, lithicpyroclastic debris generated by continued eruption of alkaline
magmas.

Stratigraphy and sedimentology of the Espinaso Formation in the Cerrillos area (Stops 2 and 3)
Daniel Larsen and Steve Taylor
Directions to Garden of the Gods (Stop 2)
From Albuquerque or Santa Fe travel 1-25 to the La Cienega exit (exit 271). Take the 1-25 east-side frontage roadsouth for 0.7 mi and turn left on NM-586. This road winds
for 6.5 mi southeastward among low hills comprised of
intrusive rocks that represent the late-stage alkaline phase of
Ortiz porphyry belt magmatism in the Cerrillos Hills area.
At the junction with NM-14, turn right and travel south for
3.8 mi. Pull off the road to the right at the "Garden of theGods" Point of Interest sign.
Geologic setting and stratigraphy of theEspinaso Formation in the Cerrillos area
The Cerrillos Hills eruptive center lies about 3 to 4 km
to the west-northwest and is composed of intrusions that
deform sedimentary strata of the Cretaceous Mancos Shaleand Mesaverde Group and Eocene Galisteo Formation(Fig. 2A). Beds of the Galisteo Formation, seen here at
Garden of the Gods, are inclined nearly vertically directly
east of the eruptive center, but dip to progressively more
shallow angles eastward. The contact between the Galisteo
and Espinaso Formations is unconformable adjacent to the
Cerrillos Hills (Disbrow and Stoll, 1957), but is transitional
to the east. The Cretaceous and lower Tertiary units in thearea are planed to a gentle slope and unconformably
overlain by late Tertiary sediments, seen immediately northof Garden of the Gods.
Strata of the Galisteo Formation represent deposition in asandy, braided alluvial system, with most of the detritus
being derived from the ancestral Nacimiento and Brazos—
Sangre de Cristo uplifts to the north (Ingersoll et al., 1990).
The transition to volcaniclastic sedimentation of the Espi-

91
naso Formation is indicated by an increase in the numberof latitic clasts, the occurrence of ash-fall tuff, and changes
in grain size and sedimentary structures in the strata.
The Espinaso Formation is composed of up to 500 m oflatitic volcaniclastic strata dominated by pebble to boulder
conglomerate and breccia with common coarse-grainedsandstone units and rare primary pyroclastic horizons. A
lava flow with limited lateral extent occurs at the base of
the section and can be seen on the low ridge 1 km to theeast. The abrupt decrease in grain size and stratigraphic
thickness of correlative intervals over a distance of 5 km
(Figs. 4, 8), the types of facies present (see descriptions for
Canada de la Cueva, Stop 3), and paleocurrent directions
indicate that these volcaniclastic strata were deposited in
an alluvial-fan setting, with the Cerrillos Hills eruptive
center as the major source of detritus.
The change in grain size and thickness of correlativeintervals with distance from the Cerrillos Hills is best exemplified by the interval with glomeroporphyritic hornblende—pyroxene-latite clasts in the lowest part of theformation. In western outcrops (section A, Fig. 4), thisinterval is 91 m thick and dominated by pebble- to boulderbearing (clasts up to 2 m across) debris-flow and sheet-flood
conglomerates and breccias. In contrast, 5 km to the south-

east (section C, Fig. 4) the interval is 25 m thick and dominated by scour-and-fill and horizontally bedded pebbly
sandstone (rare boulders up to 40 cm across).
The predominance of sheet-like debris-flow and sheetflood conglomerates and breccias suggests high-gradient,high-discharge, unconfined-flow conditions of deposition.
Bi-directional paleocurrent data from 19 scour-and-fill
structures show a strong northwest—southeast trend and asubordinate southwest—northeast trend. The paleocurrent
data combined with the facies patterns indicate generally
eastward sediment dispersal. The facies characteristics andrelationships will be discussed further at Canada de laCueva (Stop 3).
Directions to Canada de la Cueva (Stop 3)
From Garden of the Gods continue south on NM-14 for
0.9 mi to Santa Fe County Road 55A, turn left, andcontinue 2.1 mi. The log building in the side of the hill on
the left at 1.5 mi is constructed of petrified wood takenfrom the yellow pebbly sandstone in the upper part of theGalisteo Formation, a unit also found at Espinaso Ridge(Stop 1). Before the road descends into the Canada de laCueva Arroyo, turn right and park in the cleared areaadjacent to the railroad tracks.

92

Sedimentology of the Espinaso Formation in the
Canada de la Cueva area
The small-scale map of the Canada de la Cueva area in Fig.
5A shows the traverse that will be taken at this stop. The wellexposed railroad bluff section (Fig. 5B) is strati-graphicallyabove the transition zone between the Galisteo and Espinaso,
but includes the upper 14 m of the interval withglomeroporphyritic
hornblende—pyroxene-latite
clasts.
Above this interval, coarse-grained units in the section containporphyritic hornblende-latite clasts.
The section is dominated by thin (<1 m thick), discontinuous to sheet-like, pebble- to boulder-bearing, massive,
matrix-supported debris-flow breccias and vaguely stratified
sheet-flood conglomerates and breccias (Fig. 5C). The debris-flow breccias include beds with inversely graded bases,
ungraded beds, and inversely graded beds (see Walton andPalmer, 1988, for comparison). Scour-and-fill and horizontally bedded pebbly coarse-grained sandstone units are interbedded with the conglomerates and breccias. Interspersed
throughout the section are a number of thick (>1 m), laterally
continuous, pebble- to boulder-bearing debris-flow breccias.
In general, these beds show well-developed inversely gradedbases, sometimes extending well into the body of the unit,
and normally graded tops. Although evident in some of thethinner debris-flow breccias as well, the thicker breccias
commonly show signs of amalgamation of multiple-flow
units (see flows at 22 and 55 m of the section).

In some cases, pebbly coarse-grained sandstones dominate
over the conglomerate units. These intervals contain abundantscour-and-fill structures and small channels, some of which
are filled by debris-flow breccia. Paleocurrent directions
obtained from a number of scours yield northwest—southeast
trend. A larger, broader channel structure with an east—west
trend is observed at 32 m of the section; it is filled withvaguely stratified pebble to boulder conglomerate and scourand fill-bedded pebbly sandstone.
These units are interpreted as having been deposited on an
alluvial fan. The sheet-like debris-flow and sheet-flood
breccias and conglomerates represent deposition by highdischarge, unconfined flow on the fan surface. Discontinuous
conglomerate and scour-and-fill sandstones represent
deposition by more confined flow in broad, shallow channels,
as exemplified by the channel at 32 m of the section.
Primary pyroclastic deposits are rare in the Espinaso Formation and in the Cerrillos area are confined to pyroclastic
fallout. The tuff observed in Canada de la Cueva (Fig. 5A) is
only 15 cm thick and is one of the thickest observed in ourstudy area, attesting to the dearth of these types of deposits.
The tuff is comprised of mantling accretionary lapilli andlapilli-fallout tephra. Coarse-grained pyroclastic deposits
and other indications of pyroclastic and autoclastic
deposition, such as prismatically jointed clasts that are foundthroughout the Espinaso Formation (Fig. 5D) (Harlan, thisvolume), are important because they indicate that sedimentation was concurrent with volcanism.

93
Paleomagnetic constraints on the temperature of emplacement of two volcanic breccias in the
Oligocene Espinaso Formation, New Mexico: A progress report
Steve S. Harlan
Introduction

Results and discussion

Paleomagnetic analysis of the thermoremanent magnetization (TRM) of clastic deposits associated with volcanic
centers can be used to determine whether they were emplacedat elevated or "hot" temperatures (e.g. pyroclastic flows,debris flows containing hot clasts, etc.) or at "cold" or ambient
temperatures (e.g. most debris flows, glacial deposits, etc.).
Thus, paleomagnetism can be used to distinguish between
various types of unsorted and unstratified deposits commonly
found in close proximity to many vol-canoes (Aramaki andAkimoto, 1957; Hoblitt and Kellogg, 1979). More
importantly, through the use of detailed thermal
demagnetization of TRM, quantitative assessments of emplacement or "setting" temperatures can sometimes be determined (Hoblitt and Kellogg, 1979; McClelland and Druitt,1989). In this study, the paleomagnetic properties of two
breccias within the Oligocene Espinaso Formation have been
examined to provide, if possible, an estimate of the emplacement temperatures of individual clasts or of the entire
deposit. Preliminary paleomagnetic results are presented fromtwo breccias; detailed paleomagnetic and rock-magnetic
studies of these and other units are currently in progress.

Canada de la Cueva: In-situ NRM directions from Canada de la Cueva, prior to demagnetization treatment, areshown in Fig. 6a. Directions are scattered and are evenly
distributed between normal- and reverse-polarity inclinations. Of 16 samples from 13 individual blocks, 10 samples
from eight blocks (nine from prismatically jointed blocks
and one from a rounded, non-jointed clast), are characterized by a common, univectoral decay to the origin of areverse-polarity remanence during either AF or thermal demagnetization (Fig. 7a, b). The structurally corrected sample and site mean directions (D =170°, I =-35°, K = 45, α 95
= 8.4°, n = 8 blocks) are shown in Fig. 6b. The mean
direction is within 20° of the Oligocene (30 Ma) reference
direction for this locality (Irving and Irving, 1982).
Because this area appears to have experienced acounterclockwise rotation about a vertical axis and downto-the-south tilting (Brown and Golombek, 1984), directcomparison with the time-averaged Oligocene reference
direction may not be strictly appropriate. Accordingly, I
have also plotted the mean direction reported by Brown
and Golombek (1984) for the Espinaso Formation in Fig.
7b; the Canada de la Cueva result is in excellent agreementwith their observed direction.
Four other Canada de la Cueva samples that are rounded orlack obvious thermal-contraction joints contain either singleor dual-component magnetizations that, although well-defined
during demagnetization (Fig. 7c, d), have directions thatdiverge markedly from the mean direction (Fig. 6b).
Presumably, these samples contain magnetizations acquired
prior to transport and were not heated above their blocking
temperatures subsequent to deposition. A remaining sample
from a jointed block proved to be directionally unstable during
demagnetization.
The well-grouped, univectoral reverse-polarity magnetization derived primarily from the prismatically jointed clasts
and their high range of unblocking temperatures (>500°C)
suggest that the remanence is a TRM acquired during postemplacement cooling of the Canada de la Cueva deposit. The
lack of a partial TRM within the non-jointed blocks that issimilar to the mean direction derived from the jointed blocks
indicates that they were probably emplaced at temperaturesless than 300-400°C. The mixture of both hot and relatively
cold material indicates that this deposit was probably
emplaced as a debris flow that included hot blocks derived
from a nearby dome.
Arroyo del Tuerto, Espinaso Ridge: In-situ NRM directions from the Espinaso Ridge monolithologic breccia
samples are shown in Fig. 6c. Both AF and thermal demagnetization of most samples from the latite clasts and ash
matrix typically remove a low coercivity or low unblocking
temperature (<250-350°C) remanence of due northerly declination and moderate positive inclination (Fig. 7e, f, g); thisis probably a viscous remanent magnetization acquired in thepresent Earth magnetic field. At higher demagnetizations, awell-defined reverse-polarity magnetization is isolated byeither thermal or AF demagnetization in most samples (Fig.
7e, f). This magnetization is unblocked over a wide range oftemperatures (375-620°C) in individual samples,

Field and laboratory procedures
Sixteen samples from 13 individual blocks were collectedfrom a heterolithologic breccia, of probable debris-flow origin, exposed in Canada de la Cueva near Cerrillos, New
Mexico (Fig. 5). Both prismatically jointed latite blocks (Fig.
5C), presumably emplaced at elevated temperatures (Smith,
this volume), and rounded clasts were collected. At Espinaso
Ridge, in Arroyo del Tuerto, both clasts and matrix were
sampled from a nearly monolithologic breccia with an ash
matrix, which, on the basis of field relations, was probably
emplaced by a block-and-ash flow (Smith, this volume; Fig.
3). Unlike the Canada de la Cueva flow, field evidence
supporting emplacement at high temperature (such as thepresence of prismatically jointed blocks or obvious thermalcontraction joints) is largely absent in this deposit. In addition
to the latite clasts that are characteristic of this deposit,
samples from mafic inclusions and silicified or glassy accidental or accessory clasts were collected.
All samples were collected as oriented cores, using a
portable drill equipped with a nonmagnetic drill bit. Sample
NRM directions were measured using a cryogenic magnetometer at the University of New Mexico; in order to fully
characterize the rock-magnetic and directional properties ofthe samples, all samples were demagnetized by alternating
field (AF) and/or thermal methods. For statistical purposes,
only one specimen per sample has been used to calculate thesite mean directions reported below; where two or moreindividually oriented samples were collected from the same
block, their directional average was used to calculate themean direction. Characteristic directions were identified using orthogonal vector diagrams and stereographic projections. Directions from linear segments were calculated using
a least-squares technique. In general, the procedures used inthis study follow those described in detail by Kent et al.
(1981) and McClelland and Druitt (1989).

suggesting that a combination of titanomagnetite, low-Ti
magnetite, and/or hematite contribute to the remanence inindividual samples. This is confirmed by Curie temperature
determinations from representative samples that show complex, irreversible behavior with inflections at 390-400°C and520-560°C; all curves are characterized by oxidation ofmagnetite to hematite during cooling. Such complexities
probably explain the erratic behavior observed in some samples after demagnetization to high temperatures.
A small number of samples exhibits similar, but directionally distinct, moderate- and high-temperature directions;an example of such behavior is illustrated in Fig. 7g. In thissample, an initial VRM aligned close to the present Earthmagnetic field (e.g. D = 346°, I= 65°) is removed at temperatures less than 350°C. Above this temperature two distinct
linear segments, recognized by a slight inflection at 440°C, areisolated. Both the moderate-temperature (350440°C)
component (D = 143°, I= —52°) and the high-temperature(440-630°C) component (D = 120°, I = —43°) aredirectionally very similar; such behavior probably reflects aslight rotation of the clasts while they were still hot, but
subsequent to deposition. This may be due to post-emplacement compaction or deflation of the deposit.
Seven samples from three latite blocks and a mafic inclusion yield normal-polarity directions approximately antiparallel to the reverse-polarity directions observed in mostclasts (Fig. 6d). It is unlikely that this small deposit cooled
through a complete polarity transition; hence these directions
are most readily explained as a VRM that completely
dominates the remanence in these samples. Accessory or
accidental latite clasts yield either single- or dual-polarity
magnetizations that are randomly oriented (Figs. 7h, 6d).
Apparently, like many of the Canada de la Cueva clasts, theseclasts were never heated sufficiently by the Arroyo del Tuerto
flow to even partially reset their magnetizations.
Structurally corrected sample directions are shown in Fig.
6d. The mean direction, calculated from samples with wellgrouped reverse-polarity magnetizations (D = 188°, I = —48°,
K = 99.6, α 95 = 3.3°, n = 20 samples), is similar to the Oligocene reference direction (Fig. 6d), but is somewhat displaced from both the Espinaso Formation mean direction of

Golombek and Brown (1984) and the Canada de la Cueva
deposit. This difference is probably not significant given the
expected range of secular variation expected at this latitute. As
at Canada de la Cueva, this magnetization is probably a TRM
imparted during cooling of the deposit.
Maximum unblocking temperatures of 375-425°C for ashmatrix samples that contain low-Ti magnetite (Curie temperatures of 540-560°C) and the absence of a similar partial
TRM superimposed upon the accidental clasts suggest that
this deposit was emplaced at temperatures no greater thanabout 400°C. Such an emplacement temperature is consistent
with an origin as a small-volume pyroclastic flow, probably
associated with the eruption or collapse of a nearby dome
(Smith, this volume).

96
The composition of Ortiz porphyry belt extrusives as recorded in the Espinaso Formation
Gary A. Smith and Daniel W. Erskine
The magmatic centers of the Ortiz porphyry belt (Fig. 1) Using these criteria, a total of 144 clasts were sampled, 72 ofare each composed of multiple intrusions that can, in a which were point-counted and described petrographically,
general way, be divided into calc-alkaline and alkaline suites and 34 of which were analyzed for major-element compo(Disbrow and Stoll, 1957; Sun and Baldwin, 1958; Atkinson, sition. Selection of samples for modal and geochemical
1961; Maynard et al., 1990 and this volume; Coles, unpubl.). analyses was based on the degree of alteration and the order
Although there are some ambiguities in field relationship and of preference given above for similar-appearing clasts.
geochronologic data, Maynard et al. (1990) believe that calcModal compositions are presented in Fig. 8, showing
alkaline magmatism preceded alkaline magmatism, and that a stratigraphic variation in the proportion of hornblendehiatus and change in tectonic stresses intervened between among the total ferromagnesian silicates (hornblendeintrusion and extrusion of the two suites. Preserved volcanic index). Only calc-alkaline lithologies are recorded by clasts
rocks are very rare, but the composition of extruded lavas and in the sections east of Cerrillos. Stratigraphic variations
pyroclastics can be documented indirectly from study of the occur in modal mineralogy that corroborate correlations
sediments of the Espinaso Formation. This report presents the based on lithologic markers (Fig. 4). Three intervals are
results of a preliminary study of Espinaso Formation recognized. Early deposited clasts of interval I contain aconglomerate clasts.
mixture of hornblende-rich (index >0.7) and augite-rich
The temporal variation in composition of erupted (index <0.4) latites, the latter being distinctly
magmas can be approximated by careful consideration of glomeroporphyritic. The alluvial-fan sequence containing
clasts from synvolcanic conglomerates. Although any these clasts thins dramatically to the east (Fig. 8). Interval II
sediment sample will include mixtures of recently erupted contains only hornblende-rich clasts and also thins rapidly
lithologies and older material, a sampling strategy was eastward. Interval III is characterized by a wide range indesigned to emphasize the former. Samples were collected in hornblende and augite content, with repeated trends fromthe following order of preference:
hornblende-rich to augite-rich compositions.
a) Clasts containing thermal-contraction joints suggesting
At the Arroyo del Tuerto type section on Espinaso Ridge,
in-situ cooling from relatively high temperature (Fig. the abrupt decrease in hornblende index above 300 m records
5d; Harlan, this volume)the transition from calc-alkaline hornblende—augite latiteb) Clasts from monolithologic debris-flow breccias.
clasts below, to alkaline augite—biotite latite clasts in thec) Lowest occurrence of distinctive clast lithology regard- upper part of the section (Fig. 8). The discrete stratigraphic
less of depositional facies.
intervals over which detritus of the two magmatic suites arefound corroborates the temporal separation of calc-alkaline
and alkaline magmatism. Hornblende index varies greatly
within the calc-alkaline clasts. There are no obvious compositional correlations between the Arroyo del Tuerto
section and those near Cerrillos. Based on thinning trends inintervals I and II at Cerrillos, distinctive clasts associatedwith the early phases of volcanism may not be represented at
Arroyo del Tuerto, 15-20 km from vent areas, or are mixed
in the conglomerates low in the section (Fig. 3, interval I)
which, because of their polylithologic clast compositions,were not extensively sampled. Trends from hornblende-rich
to augite-rich clast compositions are apparent in the lower
200 m of the Arroyo del Tuerto section (Fig. 8) and resemble
trends in interval III near Cerrillos. The Cerrillos clasts,
however, are compositionally distinct from those at Arroyo
del Tuerto (Fig. 9).
Major-element geochemical data for Espinaso Formation
clasts from near Cerrillos and Arroyo del Tuerto are illustrated in Fig. 9, along with additional new data for clasts
from other localities, three lava flows, and three intrusions
in the Cerrillos Hills. The compositional range for Espinaso
Formation data appears representative of the magmas also
recorded in the Ortiz Mountains intrusions, based on unpublished data of David Coles (Colorado State University).
Arroyo del Tuerto subalkaline clasts are, for the most part,more siliceous than those sampled near Cerrillos. This may
account, in part, for the greater abundance of pyroclastic
debris in the Espinaso Ridge section and, along with thedissimilarity in stratigraphic variations in hornblende index
between the two localities (Fig. 8), it suggests that Espinaso
Formation detritus on Espinaso Ridge and east of Cerrillos
was derived from different sources, probably the Ortiz
Mountains and Cerrillos Hills, respectively.
More detailed consideration of clast petrology and min-eralogy, in addition to sandstone compositions, will be presented in a later paper. This study does show that valuable
information about the composition of erupted materials canbe derived from careful study of preserved volcaniclastic
sedimentary sequences in cases where the volcanic edifices
have been removed by erosion.

Stratigraphy, sedimentology, and volcanology of the Peralta Tuff Member of the Bearhead Rhyolite(Stops 4, 5, and 6)
Gary A. Smith
Directions to Stops 4, 5, and 6
Stops 4 through 6 are close to one another at, and near, aprominent feature known as the Tent Rocks (Fig. 2B).
From Albuquerque or Santa Fe, take 1-25 to exit 264 andthen head west on NM-16 for 8.3 mi. Turn right at thejunction with NM-22 and continue for 4.5 mi on NM-22 to
Cochiti Pueblo, being certain to make the marked left turn
opposite the center of Cochiti Dam. At Cochiti Pueblo, turnright onto USFS road 266 and continue 4.9 mi to Tent
Rocks. The traverse for Stop 4 follows a narrow canyon at
the northeast edge of the Tent Rocks outcrops (Fig. 2B). To
reach Stops 5 and 6, continue westward on USFS road 266
for 0.5 mi and turn right onto a narrow gravel track (Fig.
2B). To proceed farther, you will need to obtain permission
and gate keys from private-property owners.
Tent Rocks (Stop 4)
The Peralta Tuff at Tent Rocks consists mainly of conglomeratic volcanic sandstones and sandy volcanic conglomerates with intercalated primary pyroclastic deposits(Fig. 10). Nearly all of the sedimentary facies consist ofpoorly sorted scour- and fill-bedded units that are typical of
deposition by flashy, poorly confined to unconfined shallow
braided streams and sheetfloods (e.g. Smith, 1987). Massive,
generally matrix-supported debris-flow deposits are rare but
do occur. Conspicuous orange layers mark highly burrowed
and oxidized, but otherwise poorly developed paleosols,
which are typically associated with bioturbated pumice-falllayers. Most of the sediments are derived from the large,
polygenetic rhyolite-dome complex at Bearhead Peak (8.5
km to the northwest) and its peripheral pyroclasticflow fans.
Rapid downslope decrease of maximum grain size (Fig. 10)
and downslope decrease in channel depth (510 m to 2-3 mover 5 km) are characteristic of alluvial fans.
The alluvial-fan strata are intercalated with a variety ofprimary pyroclastic facies. Most of the rhyolitic pyroclastic
debris consists of angular, slightly vesicular to dense fragments whose angular morphology and limited vesiculation
are typical of phreatomagmatic eruptions (Heiken and Wohletz, 1987). These fragments were initially deposited bypyroclastic falls, flows, and surges. The fall deposits are
typically plane-bedded, with separate beds marked by differences in grain size or proportions of juvenile glassy pyroclasts and accidental lithic fragments; the bedded character
reflects variations in eruption processes at the vent. At least
20 falls occur in the section at Tent Rocks. Most are associated with paleosols and are variably bioturbated (Fig.
11B); preservation was generally limited to inactive depositional tracts on the alluvial-fan surface. Several of the finergrained ash-fall units are overlain by low- to high-angle
crossbedded ash representing eolian reworking of the fall
(Smith and Katzman, in press; see also Stop 5). The massive,
poorly sorted, white pyroclastic-flow deposits containvarying proportions of ash, pumice, dense perlitic rhyolite
glass, and darker accessory and accidental clasts of

99various compositions. None of the Peralta Tuff pyroclastic- better sorted than the surge deposits seen at Stop 4, exhibits
flow deposits are welded because of the relatively small wind-ripple and grain-flow strata, and bedsets typical of eolian
volumes and thicknesses (0.5 to 3.0 m) of the deposits and sands (Smith and Katzman, in press).
the cooling effects of the water—magma interaction during
phreatomagmatic eruptions. Some thermal oxidation did occur, resulting in orange to pink colors in some units. Pyroclastic-surge deposits are moderately sorted, crossbedded
tuffs and lapilli tuffs associated with fall beds and volcanic
bombs up to 0.5 m across.
Nearly all of the primary pyroclastic deposits occur in thelower two-thirds of Tent Rocks section (Fig. 10). The
sedimentary facies are notably coarser and darker-colored in
the upper half of the section. The grain-size change roughly
coincides with a decrease in the abundance of light-coloredvesicular and dense glassy rhyolite clasts typical of thebombs and lapilli in the primary pyroclastic units, and an
upward increase in partly to fully devitrified rhyolite clasts
derived from the interiors of rhyolite domes (Fig. 10). The
compositional and grain-size changes are interpreted asreflecting a change in sediment supply from an abundance ofrelatively fine-grained ash and lapilli, during the peak of
pyroclastic activity, to a greater contribution of coarse epiclasts generated by weathering and dissection of the dome
complexes following cessation of eruptions.
A traverse up the narrow slot canyon at the northeast
corner of the Tent Rocks crosses much of the local section
in a number of fault blocks (Fig. 11). Geologic features
along the traverse are illustrated and described in Figs. 11
and 12. The highlight of this traverse is the observation of
the tuffs of lower Peralta Canyon (Figs. 10, 11, 12), which
represent the outer flanks of an annular phreatomagmatic
tuff ring that was centered 2.5 km northwest of here (Fig.
2B). The initial phreatomagmatic explosions ejectedslurries of juvenile pyroclastics and basin-fill sediments.
The resulting debris-flow deposits are interbedded withcoarse pyroclastic-surge deposits and accretionary-lapilli
and ash-fall layers (lower unit of Fig. 12). The eruption
culminated in the accumulation of 30 m of sand-wave
crossbedded surge deposits that are notable for thepreservation of steep stoss sides, rounded to sharp crests,
and gentle lee sides of the migrating bedforms. At least 1.5
m of pumice fall accumulated on top of the surge subunit
but was largely removed by erosion of gullies and swales
on the flanks of the tuff ring. Outcrops 4 km to the northshow that this eruption followed the climax of, but
temporally overlapped with, a similar eruption at Bearhead
Peak that produced the tuffs of West Mesa (Fig. 10).
Eolian tuff, lower Peralta Canyon (Stop 5)
Stop 5 is adjacent to a rhyolite dome (Fig. 2B) that marks
the vent for the tuff ring comprising the tuffs of lower Peralta
Canyon at Tent Rocks. The margin of the dome is composedof glassy, brecciated rhyolite forming a carapace aroundflow-banded devitrified rhyolite of the dome interior that will
be seen along the road between Stops 5 and 6. This dome wascompletely buried by the alluvial fan that extended
southeastward from Bearhead Peak. Sediments exposed along
the southeast margin of the dome are stratigraphicallyequivalent to those seen at the mouth of the canyon at Stop 4.
The most remarkable feature at this locality is the occurrenceof the eroded remnant of an eolian dune that formed fromwind reworking of an ash-fall deposit upon which the eolian
tuff rests (Fig. 13). This crossbedded unit is muchJump) that extended 5 km southward from Bearhead Peak(Fig. 14). The presence of this thick lava flow on the alluvial/
pyroclastic-flow fans profoundly influenced subsequent
sedimentation, pyroclastic-flow transport and deposition, andgeometry of younger deposits. Highlights of this traverse areillustrated and described in Fig. 14.
Autoclastic breccias are well exposed along the margin of
the rhyolite of Bear Jump (Fig. 14B). These were the source
for steep debris-flow-dominated fans whose deposits are
restricted to this area (Fig. 14C) marginal to the flow, andextend less than 300 m to the east. While making thistraverse, note that structural dip is southwestward and that
depositional dips are steeply inclined (25-35°) to the east.
The most noteworthy volcanological features seen along
this traverse are associated with the tuffs of Colle Canyon.
The tuffs of Colle Canyon include as many as 25 pyroclasticflow units and less voluminous fall deposits erupted during
one eruptive episode at Bearhead Peak. The tuffs thicken
eastward where the convex-up, lobate morphology of apyroclastic-flow fan is preserved. In the area of this traverse,
some pyroclastic flows moved southward adjacent to therhyolite of Bear Jump, whereas others, having been erupted
onto the top of the lava flow close to the vent, spilled off the
The influence of lava-flow morphology on subsequent
steep front of the lava (Fig. 14). Several of these latter flow
deposition of alluvium and pyroclastic flows (Stop 6)
units contain planebedded and crossbedded intervals that can
Stop 6 consists of a traverse up a narrow tributary canyon be traced southeastward, away from the flow front, into theto Peralta Canyon that leads toward Oak Mesa and skirts the typical, massive ignimbrite. Texturally, the cross-bedded
intervals retain the poorly sorted and fine-ash-richedge of a 200 m thick rhyolite flow (rhyolite of Bear

101
character of pyroclastic-flow deposits despite illustrating
bedding that is more typical of pyroclastic-surge deposits. The
stoss-side cross-strata exhibit steep, sigmoidal forms incontrast to more gentle, tabular, lee-side strata (Fig. 14B).
Turbulence apparently resulted from acceleration of the flows
as they descended the steep flow front and decayed as theflows subsequently decelerated upon crossing lower-gra-dient slopes at the foot of the flow; there is no indication ofentrainment of air and dilution of the flows to surges as they
crossed this topographic irregularity. Other flow units
exhibit marginal onlap of the lava flow along a steep contact
near the north end of the traverse and are typical, massive
ignimbrite; this argues against generation of turbulence as amarginal boundary eddy.

Preliminary 40Ar/39Ar age determinations and paleomagnetic results from the Peralta Tuff Member,
Bearhead Rhyolite, Tent Rocks area, Jemez Mountains, New Mexico
William C. McIntosh and Steve S. Harlan
Introduction
Paleomagnetic and geochronologic studies of the Peralta
Tuff Member of the Bearhead Rhyolite and associated lavas
exposed in the Tent Rocks area of the Jemez Mountains(Fig. 10) were initiated for two purposes: (1) to use highprecision 40Ar/39Ar dating and paleomagnetic analysis toconstrain the timing and duration of the rhyolitic activity
that produced the Peralta tuff; and (2) to use paleomagnetic
analysis of thermoremanent magnetization to evaluate theemplacement temperatures of primary pyroclastic units in
the Peralta Tuff. Analysis of samples is not yet complete;
this report summarizes some preliminary findings.
40

Ar/39Ar dating

Five samples were collected for 40Ar/39Ar analysis, four
from pumice-fall deposits distributed throughout the Peralta
Tuff sequence, and a fifth from the rhyolite of lower Peralta
Canyon (Fig. 10). The fallout-deposit samples were split
into pumice and loose crystal fractions, and sanidine separates (>99.5% purity) were prepared from these splits andfrom the rhyolite lava. To date five samples have been analyzed at the U.S. Geological Survey isotopic dating facility
in Reston, Virginia, using analytical procedures similar to
those described in McIntosh et al. (in press).
Results from samples of the stratigraphically lower pumicefall deposit, the tuff of West Mesa, and the rhyolite lava show
generally flat 40Ar/39Ar age spectra (Fig. 15), yielding
statistically indistinguishable plateau ages ranging from 6.81 ±0.2 to 6.87 ± .02 Ma (Table 1). The remaining two samples,from pumice falls immediately below the tuffs of Canada
Camada (Fig. 10), failed to yield age plateaus (Fig. 15),
probably due to the presence of abundant large glass

inclusions. Their total gas ages (6.71 and 7.08 Ma, Table 1)
bracket the plateau ages from the samples lower in thesequence. Although total gas ages are considered to be lessreliable than the plateau ages, these preliminary data areconsistent with eruption of the entire Peralta sequence within abrief interval of time.

102
Paleomagnetism
Oriented samples for paleomagnetic analysis were collectedfrom 25 sites throughout the Peralta Tuff section shownin Fig. 10. At three sites in lavas, 8 to 12 cores werecollected using a portable drill; individual cores were oriented
using both sun and magnetic compasses. At 22 sites
in tuffs, 8 to 12 pumice or lithic clasts were collected using
either the drill or taking oriented block samples. Samples fromselected sites have been analyzed using a cryogenic
magnetometer at the University of New Mexico. Initially,
pilot samples from each site were subjected to stepwise
alternating field (AF) demagnetization to peak fields of 140
mT. For three sites, thermal demagnetization behavior has also
been investigated using stepwise heating to temperatures of
600°C.
Samples from the three sites in the Bearhead Rhyolite lavas
are characterized by nearly univectoral behavior during AF demagnetization, indicating the presence of a single
well-defined magnetization component throughout the demagnetization interval (Fig. 16a). Sample remanence directions are well-grouped at the site level; all three sites yieldreverse-polarity site mean directions with southerly
declinations (D = 130° to 180°) and moderate negative inclinations (I = —45° to —60°).
Demagnetization behavior of samples from sites in the tuffs
is usually more complex and typically varies with lithology. White, highly inflated pumice clasts typically yield
low-intensity, randomly oriented remanence directions andunstable AF demagnetization behavior, probably reflecting
Fe—Ti oxides in the superparamagnetic size range. Denser
vesiculated rhyolite and accidental or accessory lithic clasts
show natural remanent magnetizations that tend to be poorly
grouped at the site level, although reverse-polarity directions
are most common. Most of the samples show approximately
univectoral decay of the magnetization vector during AF
demagnetization (Fig. 16b), but more complex AF demagnetization behavior was observed in some samples (Fig.
16c).
Preliminary results from the three sites that were thermallydemagnetized indicate that samples from these ignimbrites carry a paleomagnetic record of their emplacementor "setting" temperatures. Magnetizations unblocked at low
temperatures (250-440°C) are generally well-grouped anddisplay reverse-polarity directions, whereas magnetizations
unblocked at higher temperatures tend to be randomly directed(Fig. 16d). The lower-temperature magnetizations areinterpreted to have been acquired by cooling subsequent todeposition of the clasts, thus recording emplacement at temperatures between 250 and 400°C. This method of paleomagnetically estimating emplacement temperature has beensuccessfully applied elsewhere (e.g. Hoblitt and Kellogg,
1979).
Although the paleomagnetic study is still in its initialstages, the poorly grouped but generally reversely polarized
directions shown by well-behaved samples from all sites in the
Peralta Tuff are interpreted as resulting from emplacement
below their maximum blocking temperatures (e.g. 580°C, the they were emplaced during a single geomagnetic reversedCurie temperature of pure magnetite) during a period of polarity interval near 6.85 Ma. The geomagnetic polarity
time scale of Berggren et al. (1985) shows three potentiallyreversed geomagnetic polarity.
correlative reverse-polarity intervals at 6.50 to 6.70 Ma, 6.78
to 6.85 Ma, and 7.28 to 7.35 Ma. The maximum eruptive
Discussioninterval for the Peralta Tuff is therefore estimated at 70 toPreliminary data from the Peralta Tuff and Bearhead Rhy- 200 Ka. Additional 40Ar/39Ar work now in progress will
olite in the Bearhead Peak—Tent Rocks area indicate thatfurther constrain this eruptive interval.

103
Combined references
Aramaki, S., and Akimoto, S., 1957, Temperature estimation of
pyroclastic deposits by natural remanent magnetism: American Journalof Science, v. 255, pp. 617-627.
Atkinson, W. W., Jr., 1961, Geology of the San Pedro Mountains, SantaFe County, New Mexico: New Mexico Bureau of Mines & Mineral
Resources, Bulletin 77, 50 pp.
Bachman, G. 0., 1975, Geologic map of the Madrid quadrangle, Santa Fe
and Sandoval Counties, New Mexico: U.S. Geological Survey, Geologic Quadrangle Map GQ-1268, scale 1:62,500.
Baldridge, W. S., Damon, P. E., Shafiqullah, M., and Bridwell, R. J.,
1980, Evolution of the central Rio Grande rift, New Mexico: New
potassium-argon ages: Earth and Planetary Science Letters, v. 51, pp.
309-321.
Berggren, W. A., Kent, D. V., and Flynn, J. J., 1985, Jurassic to Paleogene: Part 2, Paleogene geochronology and chronostratigraphy; in
Snelling, N. J. (ed.), The chronology of the geological record:
Geological Society of London, Memoirs, v. 10, pp. 141-195.
Brown, L. L., and Golombek, M. P., 1984, Block rotations in the Rio
Grande rift, New Mexico: Tectonics, v. 5, pp. 423-438.
Disbrow, A. E., and Stoll, W. C., 1957, Geology of the Cerrillos area,
Santa Fe County, New Mexico: New Mexico Bureau of Mines &Mineral Resources, Bulletin 48, 73 pp.
Gardner, J. N., Goff, F., Garcia, S., and Hagan, R. C., 1986,
Stratigraphic relations and lithologic variations in the Jemez Volcanic
Field, New Mexico: Journal of Geophysical Research, v. 91, no. B2,
pp. 17631778.
Heiken, G., and Wohletz, K., 1987, Tephra deposits associated with silicic
domes and lava flows; in Fink, J. H. (ed.), The emplacement of silicic
domes and lava flows: Geological Society of America, Special Paper
212, pp. 55-76.
Hoblitt, R. P., and Kellogg, K. S., 1979, Emplacement temperature ofunsorted and unstratified deposits of volcanic rock debris as determinedby paleomagnetic techniques: Geological Society of America, Bulletin,v. 90, pp. 633-642.
Ingersoll, R. V., Cavazza, W., Baldridge, W. S., and Shafiqullah, M.,
1990, Cenozoic sedimentation and paleotectonics of north-central New
Mexico: Implications for initiation and evolution of the Rio Grande rift:
Geological Society of America, Bulletin, v. 102, pp. 1280-1296.
Irving, E., and Irving, G. A., 1982, Apparent polar wander paths, Carboniferous through Cenozoic and the assembly of Gondwana: Geophysical Surveys, v. 5, pp. 141-188.
Kautz, P. F., Ingersoll, R. V., Baldridge, W. S., Damon, P. E., andShafiqullah, M., 1981, Geology of the Espinaso Formation (Oligocene),
north-central New Mexico: Geological Society of America, Bulletin, v.
92, pt. I, pp. 980-983, pt. II, pp. 2318-2400.

Kent, D. V., Dragoslav, N., Pescatore, T., and Sparks, R. S. J., 1981,
Paleomagnetic determination of emplacement temperature of Vesuvius
AD 79 pyroclastic deposits: Nature, v. 290, pp. 393-396.
Lucas, S. G., 1982, Vertebrate paleontology, stratigraphy, and biostratigraphy of Eocene Galisteo Formation, north-central New Mexico: New
Mexico Bureau of Mines & Mineral Resources, Circular 186, 34 pp.
Maynard, S. R., Martin, K. W., Nelsen, C. J., and Schutz, J. L., 1990,
Geology and gold mineralization of the Ortiz Mountains, Santa Fe
County, New Mexico: Mining Engineering, v. 42, pp. 1007-1011.
MacDonald, G. A., and Katsura, T., 1964, Chemical composition of Hawaiian lavas: Journal of Petrology, v. 5, pp. 82-133.
McClelland, E. A., and Druitt, T. H., 1989, Paleomagnetic estimates of
emplacement temperatures of pyroclastic deposits on Santorini, Greece:
Bulletin of Volcanology, v. 51, pp. 16-27.
McIntosh, W. C., Sutter, J. F., Chapin, C. E., and Kedzie, L. L. (in press),
High-precision 40Ar/39Ar sanidine geochronology of ignimbrites in the
Mogollon-Datil volcanic field, southwestern New Mexico: Bulletin ofVolcanology.
Smith, G. A., 1987, Sedimentology of volcanism-induced aggradation in
fluvial basins: Examples from the Pacific Northwest, U.S.A.; in Ethridge, F. G., Flores, R. M., and Harvey, M. D. (eds.), Recent developments in fluvial sedimentology: Society of Economic Paleontologists
and Mineralogists, Special Publication 39, pp. 217-229.
Smith, G. A. (in press), Facies sequences and geometries in continental
volcaniclastic sediments; in Fisher, R. V., and Smith, G. A. (eds.),
Sedimentation in volcanic settings: SEPM, The Society for Sedimentary
Geology, Special Publication.
Smith, G. A., and Katzman, D. (in press), Discrimination of eolian andpyroclastic surge processes in the generation of cross-bedded tuffs,
Jemez Mountains volcanic field, New Mexico: Geology.
Steams, C. E., 1953, Early Tertiary volcanism in the Galisteo-Tongue
area, north-central New Mexico: American Journal of Science, v. 251,
pp. 415-452.
Sun, M.-S., and Baldwin, B., 1958, Volcanic rocks of the Cienega area,
Santa Fe County, New Mexico: New Mexico Bureau of Mines &Mineral Resources, Bulletin 54, 80 pp.
Walton, A. W., and Palmer, B. A., 1988, Lahar facies of the Mount Dutton
Formation (Oligocene-Miocene) in the Marysvale Volcanic Field,
southwestern Utah: Geological Society of America, Bulletin, v. 100, pp.
1078-1091.
Waresback, D. B., and Turbeville, B. N., 1990, Evolution of a PlioPleistocene volcanogenic-alluvial fan: The Puye Formation, Jemez
Mountains, New Mexico: Geological Society of America, Bulletin, v.
102, pp. 298-314.

New Mexico Bureau of Mines & Mineral Resources Bulletin 137, April 1991

105

Quaternary and Neogene landscape evolution: A transect across the
Colorado Plateau and Basin and Range provinces in westcentral and central New Mexico
John W. Hawley and David W. Love
New Mexico Bureau of Mines & Mineral Resources, Socorro, New Mexico 87801

with supplemental articles by
J. L. Betancourt, R. M. Turner, and S. Tharnstrom / L. Thompson, E. Mosley-Thompson, J. L. Betancourt, D. W. Love, A. Wilson, G. Leonard,
and R. S. Anderson / P. G. Drake, C. D. Harrington, S. G. Wells, F. V. Perry, and A. W. Laughlin / R. P. Lozinsky, J. W. Hawley, andD. W. Love / W. C. Haneberg / R. C. Lohmann, J. M. Davis, D. W. Love, and F. Phillips / B. D. Allen / and P. J. Slavin

Introduction
Emphasis of this transect across west-central and central
Acknowledgments
New Mexico is on the Quaternary to late Miocene
The compilers of this field guide thank the many persons
evolution of the great variety of landforms and associated and organizations who have assisted in this undertaking.
surficial deposits that, in aggregate, make up the landscape First and foremost are all the contributors. Particular thanks
in parts of two contrasting physiographic provinces: the goes to Steve Wells, who organized the field trip, developed
Colorado Plateau (Acoma—Zuni and Navajo sections), and the outline, and solicited most of the contributed articles.
the Basin and Range (Mexican Highland and Sacramento This field guide could not have been compiled without thesections). Grants, at the base of the Pliocene Mount Taylor strong support of Frank E. Kottlowski, Director and State
volcanic center, is the starting point for Day 1 and Day 2 Geologist, and Jiri Zidek, Chief Editor—Geologist, New
tours of the upper Rio San Jose, and Puerco and Zuni River Mexico Bureau of Mines & Mineral Resources, and otherbasins (Fig. 1). The latter two drainage systems are west of members of the Bureau scientific and technical staff. The
the Continental Divide in the Gallup and Zuni Pueblo area; New Mexico Geological Society, Inc. graciously allowed
and all three head in the Zuni uplift.
citation and use of illustrations from their Guidebook and
The second-day tour between Grants and Mountainair Special Publication series.
crosses the Albuquerque basin of the Rio Grande rift, as well
Persons and organizations who for this field trip, or in theas the flanking Lucero and Manzano—Los Pinos uplifts at past, have granted us access to lands that they own or
the margins of the Colorado Plateau and Great Plains administrate are also gratefully acknowledged. They
structural provinces (Fig. 2). Highlights of the first two days include the Laguna and Zuni Tribal Governments, theinclude Quaternary alluvial chronologies and changing Navajo Nation, the U.S. Forest Service, the U.S. Bureau ofenvironments, both in the Puerco and Zuni basins of the Land Management, the U.S. Fish and Wildlife Service, theColorado River system and in the Rio San Jose and Rio New Mexico State Land Office, and the Huning Land
Puerco basins of the upper Rio Grande. Pliocene and Qua- Trust. Individual articles by field-guide contributors also
ternary volcanic and neotectonic features of Zuni—Bandera, contain acknowledgment sections.
Mount Taylor, and Lucero fields are also covered.
Part two of the second-day tour includes an introduction to
ongoing studies of late Quaternary lacustrine, playa, eolian, Directions and comments en-route from Albuquerqueto Grants
and paleohydrologic features of the closed-drainage Estancia
basin east of Mountainair (Fig. 2) and in the transition zone
Take 1-40 west from its interchange with 1-25 (just east
(Sacramento section) between the Basin and Range and Great of MP 159). Proceed 73 mi to Grants. MP 86. Prepare to
Plains physiographic provinces. The highest shoreline take Exit 85 in 0.5 mi to motel area, and access todeposits of late Wisconsin Lake Estancia are visited on the Business-40 through Grants and Mount Taylor area. The
morning of Day 3; and after traversing the Rio Grande rift- following synopsis of geologic features along the way isbounding Sandia—Manzano uplifts, via Tijeras Canyon, the taken from the New Mexico Geological Society 40th Field
field trip concludes with a tour of the Albuquerque—Rio Conference road log by Maxwell et al. (1989).
This segment of the tour takes us west 75 mi from AlRancho urban area. In addition to urban environmental geology, tour stops emphasize Miocene to Pleistocene basin-fill buquerque to Grants, New Mexico. The route crosses the Rio
stratigraphy, Quaternary evolution of the Rio Grande valley, Grande valley (to just west of MP 156) and ascends across
two major Pleistocene river terraces to an early middle
and Plio-Pleistocene basaltic volcanism and neotectonism.
The New Mexico Highway Geology Map (1982) pub- Pleistocene geomorphic surface, the Llano de Albuquerquelished by the New Mexico Geological Society, Inc. (MP 150-145), that marks the culmination of basin filling in(Campus Station, Socorro, New Mexico 87801) is the Albuquerque basin. The Llano surface has been faulted
recommended as a supplement to this field guide. and southward-trending valleys have been cut as much as
Stratigraphic nomenclature for the region (Anderson, Lucas, 300 m into the basin fill (Santa Fe Group) by both the Rio
Grande and Rio Puerco. The 150 ka Albuquerqueand Love, 1989) is summarized in Fig. 3.

109
Volcanoes erupted on the eastern margin of the Llano andflowed to terraces along the Rio Grande (Stop A-4).
West of the Rio Puerco (near MP 140) the route crosses
the Rio Puerco fault zone, which forms the structural boundary between the Colorado Plateau and the Basin and Range
province (Rio Grande rift). Adjacent to 1-40, the 620,000 yrs
old Lava Creek B ash is preserved in sediments filling theApache graben, a relatively small structure within the fault
zone. On the Colorado Plateau, the route passes nearly flatlying multicolored Triassic, Jurassic, and Cretaceous

continental and marine strata and Plio-Pleistocene volcanoes
and flows. West of Exit 126 and the junction with NM-6, theroute follows the Rio San Jose, a major tributary of the Rio
Puerco. Its valley has been affected repeatedly during
Pliocene-Quaternary time by large-scale erosion and backfill processes, as well as by episodic emplacement of basalt
flows (discussed at Stops 1-7 and 2-1). This segment of thetrip to Grants follows the eastbound route covered in Part 1of the Day 2 section of the tour guide.

Day 1: Grants-Gallup-Zuni Mountain loop

SummaryThis segment of the tour traverses parts of the AcomaZuni
and Navajo sections of the Colorado Plateau province
between Grants, Gallup, and Zuni Pueblo (Fig. 1). Emphasis
is on (1) the Quaternary and late Neogene geomorphic
evolution of the upper Little Colorado (Puerco-Zuni) Riverbasin of west-central New Mexico; and (2) Quaternary environments of the Zuni-Bandera volcanic field which straddles the Continental Divide to the upper valley of the Rio
San Jose near Grants. Tour highlights include discussions of
Quaternary history and late Holocene ecological changes on
the summit and flanks of the Zuni Mountains (Stops 11, 1-2,
1-5, and 1-6); mid-to-late Quaternary alluvial chronologies
in the upper Puerco and Zuni River valleys west of theContinental Divide (Stops 1-2, 1-3, and 1-5); Pleistocene
basaltic volcanism of the Zuni-Grants area (Stops 1-5, 1-6,
and 1-7); and Mio-Pliocene erosion surfaces and associatedalluvial deposits (Bidahochi Formation) on the western flankof the Zuni uplift (Stop 1-4).
Most of the comments en-route to the tour stops are adaptedfrom road logs by Anderson et al. (1989); Anderson, Love, andLucas (1989); Chamberlin et al. (1989); Colpitts et al. (1989);
Maxwell et al. (1989); Hawley, Crumpler, and Wells (1982);
and Wells et al. (1983). Supplemental geologic maps
recommended for this tour are by Dillinger (1990), Goddard(1966), Hackman and Olson (1977), and Maxwell (1986).
Directions and comments en-route to Stop 1-1
From the parking lot of the Inn at Grants go south 0.5 mi
to 1-40 Interchange (85) and take westbound onramp to 140. Milepost (MP) 85 is the 0.0 mi point. Proceed 51 minorthwest on 1-40 to Exit 33 (McGaffey, Fort Wingate),and 10 mi south on NM-400 to McGaffey and Stop 1-1(optional excursion). Optional Stop 1-2 is 7 mi north of
McGaffey and overlooks Fort Wingate east of MP 7 on
NM-400. The tour route from Grants to Fort Wingate isalong the northern edge of the Acoma-Zuni section of theColorado Plateau physiographic province (Hawley, 1986a).
0.0 mi, MP 85, crossing basalt flow from El Calderon
center for next 1 mi. The vent is located about 16 mi (26 km)
to the southwest in the Bandera-Oso Ridge area viewed at
Stop 1-6. Note nearly continuous eolian veneer and locallywell-developed calcic soils in cuts. Preliminary attempts at
K-Ar and magnetostratigraphic dating of this flow haveyielded ages ranging from early to late Pleistocene. MP 84,

route for next 3 mi on upper Pleistocene aa-lava flow fromthe Paxton Springs center. The final part of today's tour
(Stops 1-6, 1-7) follows this flow from its source area inthe southeastern Zuni Mountains through Zuni Canyon.
MP 83, at 9:30 in foreground spring-travertine moundon Chinle Formation is surrounded by Paxton Springs
flow. Beyond the mound at the foot of Zuni Mountains
near San Rafael are active springs described by White(1989). The dip-slope flanking the northeastern Zuni upliftfrom 9:00 to 12:00 is capped by limestones of the Permian
San Andres Formation. The Precambrian-cored Zuni uplift
forms the southern edge of the San Juan basin and has been
a positive area during most of Phanerozoic time.
Chamberlin and Anderson (1989a) describe a model for thecomplex, horizontally directed, compressive forces that
may have produced the major structural components of theZuni uplift during the late Laramide (Eocene).
2.9 mi, Exit 81 to right; access to West Grants and Zuni
Canyon Rd. (0.2 mi south of interchange); and (via NM-53)
to San Rafael, Bandera, and Oso Ridge craters, El Morro
National Monument, and Zuni Pueblo. At 1:00 to 3:00, upper
Pliocene basalt of Black Mesa (southwest end of West Grants
Ridge) is flanked by landslide deposits on Triassic Chinle
Formation mudstones and sandstones. 4.0 mi, mouth of Zuni
Canyon at 10:30; crossing Rio San Jose at north edge of
Paxton Springs flow and south edge of Milan. Ditch ahead onright (4.3 mi) exposes older (unidentified) flow with calcic
soils partly formed in capping deposits.
5.6 mi, Exit 79 to right; access via NM-605 to Milan,
San Mateo, and Ambrosia Lake areas of the Grants uranium
region (Kelley, 1963; Rautman, 1980; Brod and Stone,1981) and Lee Ranch coal mine. 6.8 mi, "Thoreau 25,
Gallup 56" mi; at 2:30, about 4 mi to the northeast, the large
white tailings pile of United-Nuclear-Homestake Partners
mill is visible on the broad floor of San Mateo Creek valley.
See Reith and Thomson (1991) for discussion of UMTRA
program to stabilize large uranium-mill tailings in the San
Juan basin region.
6.0 mi, MP 76, at 10:00 to 12:00 limestones and sandstones of the San Andres and underlying Glorieta Formations
form cuesta dip-slope on the northeastern flank of the Zuni
Mountains. Underlying gypsum, sandstone, and limestone of
the Lower Permian Yeso Formation and basal Abo Formation redbeds overlap an undulating erosion surface on
Precambrian metamorphics (see Colpitts et al. , 1989, and110road log between Stops 1-6 and 1-7). MP 77, crossing
cuesta dip-slope for next 2 mi, with excellent roadcut
exposures of San Andres limestones.
10.0 mi, MP 75, Haystack Mountain at 2:00 is the site of
the initial (1950) Grants region uranium discovery in theJurassic Todilto Formation (Kelley, 1963). This butte iscapped by Upper Cretaceous Dakota Sandstone. Jurassic
rocks exposed in the south face are, in ascending order, theEntrada Sandstone (massive red), Todilto Formation (graylimestone and gypsum), Wanakah Formation (reddishbrown sandstone and mudstone), and members of theMorrison Formation (variegated shales, siltstones, andsandstones). Triassic Chinle and Moenkopi Formations(mudstones and sandstones) underlie the broad strike valley
between Haystack Mountain and this point. El Tintero coneon the valley floor just south of the mountain is the sourceof "the Malpais" basalt flow (middle? Pleistocene) thatcovers much of the floor of this part of the upper Rio San
Jose (and Blue-water Creek) valley. Fine- to mediumgrained alluvial valley fill (mid-to-late Quaternary) is
locally thick (as much as 3040 m, 100-130 ft; Gordon,
1961) and commonly has expansive (vertic) surface soils.
11.5 mi, Anaconda (uranium) mill site to right in middle
distance. 12.5 mi, Exit 72 to Bluewater. Small outcrops androadcut exposures of El Tintero flow for next 4.0 mi. 14.5
mi, mouth of Bluewater Canyon at 10:00 is cut in San
Andres limestones and sandstones. Lower Chinle/Moenkopi
redbeds crop out to right of canyon and are capped bySonsela Sandstone Bed of the Petrified Forest Member ofChinle Formation. Roadcuts for the next 46 mi (to McGaffey-Ft. Wingate area between Stops 1-1 and 1-2) areprimarily in the Petrified Forest Member; and the Sonsela
Sandstone Bed is the main cuesta dip-slope caprock (Hackman and Olson, 1977).
15.5 mi, Sonsela Sandstone caps dip-slope remnant toleft. For next 20 mi the route follows strike valley of upper
Rio San Jose developed in middle and upper part of the
Chinle Formation. 16.0 mi, Cibola-McKinley County line.
El Tintero cone at 2:00 and roadcuts ahead in El Tintero
basalt. 17.0 mi, MP 68, crossing western edge of El Tintero
flow. Haystack Mountain (7833 ft, 2387 m) at 2:30.
19.0 mi, rising on Sonsela Sandstone dip-slope; roadcuts
for the next 34 mi primarily in Sonsela Sandstone Bed andcontiguous mudstones and sandstones of the Petrified Forest
Member. 20.4 mi, at 3:00, across valley, Morrison Formation (variegated to grayish-brown slopes) is downthrown
to the east against Entrada Sandstone (red cliffs) by northtrending Bluewater fault. 21.5 mi, Exit 63 to Prewitt.
24.0 mi, MP 61, Escalante Generating Station of Plains
Electric Power Coop at 3:00, at base of Entrada Sandstone
cliffs, has an initial capacity of 233 megawatts, which isfurnished to 16 rural cooperatives. Coal used in theoperation comes by unit trains from Lee Ranch mine northof Ambrosia Lake.
29.0 mi, MP 56, from 2:00 to 3:00 a Jurassic-Cretaceous
stratigraphic sequence (Entrada, Todilto-Wanakah, Morrison, capping Dakota Formation) is exposed in prominent
cuesta scarp north of (strike) valley of upper Rio San Jose.
The sequence is capped by a mid-Tertiary diabase sill at
Mount Powell (2:00) on Continental Divide. 31.2 mi, Exit
53 at Thoreau. Access, via NM-371, to Crownpoint andChaco Canyon area of the San Juan basin and the site of the1983 American Geomorphological Field Group Conference(Wells et al., 1983). 33.0 mi, MP 52, at 3:00, relay

towers on Mount Powell (8748 ft, 2666 m) are on Continental Divide. Route on "spectacular" ascent to Continental
Divide for next 3.8 mi.
36.8 mi, Exit 47. Continental Divide (about 7270 ft,
2216 m). Access to "Top 0' The World" service center.
Route crosses from upper basin of the Rio San Jose-Rio
Puerco system (tributary to the Rio Grande) to the
Colorado River watershed and the upper end of the Puerco
River (Little Colorado) drainage basin. 37.0 mi, MP 48,
strike valley on Chinle Formation dip-slope continues tothe west another 20 mi almost to The Hogback (Nutria
monocline) belt at the east edge of Gallup.
40.1 mi, Exit 44 to Coolidge. 44.9 mi, Exit 39 to Rest
Area and refinery. 46.7 mi, valley-border surfaces to righthave thin cap of gravelly alluvium (mid? to late
Pleistocene). 48.0 mi, Exit to Iyanbito. 49.0 mi, MP 36,
shallow channel to right is South Fork of the Puerco River(Rio Puerco of the West in older literature). 50.0 mi, MP
35, thin Pleistocene terrace gravels cap Sonsela Sandstone
outliers ahead on right.
51.0 mi, MP 34, prepare to exit 1-40 in 0.5 mi. 51.5 mi,take Exit 33 to McGaffey, Fort Wingate, and optional excursion to Stops 1-1 and 1-2 (via NM-400). 51.2 mi, turnleft (west) on frontage road. 51.5 mi, turn left (south) on
NM-400; cross 1-40 and shallow channel of South Fork of
the Puerco River (51.8 mi).
52.0 mi, MP 10 on NM-400, route ahead crosses low
strath terrace on Petrified Forest Member of Chinle Formation capped by thin gravelly fill and then ascends tosummit of cuesta dip-slope on Sonsela Sandstone Bed. 53.0
mi, MP 9, descending cuesta-scarp slope cut in sandy mudstones of the Petrified Forest Member. 53.9 mi, Fort Wingate High School to right. 54.3 mi, historical marker to left
at entrance to 19th-century military post of Fort Wingate(see Stop 1-2 discussion). 54.8 mi, Saint Elenors Church toright. Winding road and steep grade ahead, with roadcuts
in mudstones and limestone-pebble conglomerate bed in
Chinle to left.
55.0 mi, MP 7, parking area on left for Fort Wingate
overlook and Stop 1-2 (optional); proceed through cattle
guard (mi 55.1) on NM-400. Note typical development ofpiñon-juniper forest (Pinus edulis and Juniperus monosperma) ahead. Winding road and steep grade (55.4 mi)with roadcuts in lower Chinle mudstones and sandstones.
Landslides are common.
56.0 mi, MP 6, note scattered ponderosa pine (Pinus
ponderosa) in piñon-juniper forest. Varicolored Petrified
Forest Member crops out on both sides of road. 56.3 mi,crest of grade; extensive ponderosa stands in forest ahead(56.7 mi); mottled strata are characteristic of the basal
Chinle Formation.
57.0 mi, MP 5, cattle guard; entering Cibola National
Forest. 57.4 mi, San Andres Formation thins in this areaand may be locally absent, with the base of the Triassic
redbed sequence resting directly on Glorieta Sandstone.
57.6 mi, Glorieta Sandstone crops out in canyon wall toright, and roadcut ahead on left (57.9 mi) exposes San
Andres/ Glorieta (limestone/sandstone) contact.
58.0 mi, MP 4, approaching head of Milk Ranch Canyon
cut in San Andres Formation limestones, here as much as 50
ft (16 m) thick. 59.0 mi, MP 3, on summit of northwestern
Zuni Mountains, with high ridges capped with thick fluvial
sandstones of the Chinle Formation (probably Sonsela
Sandstone Bed). 59.4 mi, cattle guard. Quaking Aspen

111
Campground ahead on right. 61.5 mi, drainage divide atabout 7805 ft (2379 m). Forest Service road to right crosses
steep western slope of Zuni Mountains and continues to
NM-602 (Stop 1-4 area) via "west wall" of the Nutria
mono-cline and Bread Springs (see Anderson, Love, andLucas, 1989).
62.0 mi, MP 0, at McGaffey (elevation 7780 ft, 2371
m). Area of Stop 1-1.
Stop 1-1 (optional): Forest ecosystems of Zuni Mountainarea: Impact of timber management practices. Optional
discussion of late Holocene—historical changes in forest ecosystem by Loren Potter and S. G. Wells.
Retrace route on NM-400 to overlook at Fort Wingate
Centennial Monument (MP 7) and Stop 1-2.
Stop 1-2 (optional): Western Puerco River basin from
Fort Wingate Centennial Monument (J. W. Hawley and
D. W. Love). This stop (elevation about 7140 ft, 2176 m)
provides an excellent view of the upper Puerco River basin
and the southern border of the Navajo section of the Colorado Plateau (Fig. 1-1). The Acoma—Zuni section to thesouth (Hawley, 1986a) was originally included in the northern Datil section (Fenneman, 1931; Hunt, 1974) which comprised a poorly defined, transitional physiographic
subdivision between the southeast Colorado Plateau andBasin and Range provinces.
Fort Wingate was originally established as Fort Fauntleroy in 1860, but the post's garrison withdrew in 1861 because of the Confederate invasion of New Mexico's Rio
Grande valley. Only a mail station was maintained here
during the Civil War. In 1868 the post was re-establishedas Fort Wingate (for Capt. Benjamin Wingate, who died ofwounds received in the Battle of Valverde south of Socorro
in 1862).
The prominent cliffs north of 1-40, on the opposite sideof South Fork of Puerco River valley, expose the following

Mesozoic sequence (ascending): Chinle (Triassic), red Entrada Sandstone (basal Jurassic unit), thin white band of
Todilto limestone and gypsum, Beclabito and Horse Mesa
Members of Wanakah Formation, light-colored Cow Springs
Sandstone and variegated Morrison Formation (uppermost
Jurassic unit), and thin, generally cliff-forming Dakota Formation capping the section. The Navajo Church (ChurchRock) and Pyramid Rock pinnacles and "White Cliffs" westof the mainstem Puerco River valley (Fig. 1), are developedin Westwater Canyon Member of Morrison Formation.
Retrace route to 1-40 Interchange 33. Turn left on northfrontage road and proceed 0.1 mi to 1-40 West access road.
Directions and comments en-route to Stop 1-3
0.0 mi, turn left onto 1-40 west. 0.5 mi, MP 33, proceed
5.5 mi west down the valley of Puerco River's South Fork
to Exit 26 and Stop 1-3a at The Hogback in eastern
Gallup. 1.5 mi, MP 32, note large dunes at base of Entrada
Sandstone and underlying Owl Rock Member of Chinle
Formation. 3.0 mi, at bridge over railroad spur, bunkers ofFort Wingate U.S. Army Ordnance Depot to left; Pyramid
Rock and Navajo Church (Church Rock) on skyline at
2:00 and 3:00. 3.5 mi, MP 30, crossing South Fork of the
Puerco River; roadcuts in Chinle Formation mudstone andsandstone. 4.1 mi, Entrada Sandstone cliffs at Redrock
State Park to right. Site of Navajo—Intertribal Ceremonial
event every August. Note large dunes at cliff base. 5.5 mi,
MP 28, at 1:00 to 2:00 Dakota-capped light-colored
Morrison Formation/Cow Springs Sandstone escarpment
forms the "White Cliffs." Dakota Sandstone andMesaverde Group-capped cuestas of the northwestern
Zuni uplift are on the skyline.
6.5 mi, MP 27, prepare to take next exit (Gallup East).
Confluence of mainstem and South Fork of the Puerco River
on valley floor to right. 7.1 mi, take Exit 26 to right. 7.5 mi,turn left on US-66 (Business 1-40) and proceed west throughunderpass into Gallup. 8.0 mi, watergap through double
hogback belt of the west-dipping Nutria monocline ahead(Dakota Sandstone to the east, strike valley in Mancos Shaleand Gallup Sandstone to the west). This steep monoclinal
structure, "The Hogback," marks the western edge of theZuni uplift and eastern border of the Gallup sag. Prepare forright turn on a narrow road in 0.5 mi.
8.5 mi, turn right on graded road to Superman Canyon.
Continue under ATSF Railroad and across bridge overPuerco River. 8.55 mi, park off road at eastern base of
Dakota Sandstone hogback; walk to tip of hogback for Stop
1-3 orientation and then continue down Puerco valley floor
about 200 m to Stop 1-3a.
Stop 1-3a: Alluvial fills of the Puerco River valley—Gallup
Hogback area, and impacts of the 1979 uranium tailings
spill north of Church Rock (D. W. Love and J. W. Hawley).
The Puerco River of the West (as opposed to the Rio Puercoof the East and Rio Puerco of the North, both east of theContinental Divide) drains about 558 me (1445 km') east of
Gallup. Its drainage area expands to several thousand square
miles before it joins the Little Colorado River near Holbrook,
Arizona, 90 mi (145 km) to the west. The maximum historic
flood was 12,000 ft3/s (340 m3/s) on July 17, 1972. Stage was
15.3 ft (4.66 m). Over a million dollars of damage was done in
Gallup. Floods in Gallup have repeated since then, but the
Army Corps of Engineers is planning a flood-control channel
north of the natural channel.

112
In 1979 a uranium-mill tailings dam about 14 mi (23km) upstream from this stop breached and 1.1 ton (1 x 10^3kg) of radioactive and acidic sediment and 124 x 106 ft3(3.5 x 106 m3) of liquid wastes swept down the Puerco
River (Gallaher and Goad, 1981; Miller and Wells, 1986).
Miller studied the downstream variation in conditions for
sediment storage along the Puerco River and looked forcontaminated sediments five years after the spill. He wasunable to locate consistent storage areas of contaminatedsediments. He attributed the lack of contaminatedsediments to dilution and possible eolian transport out of
the arroyo system. Gray et al. (in progress) have foundample evidence of ground-water contamination beneath the
Puerco River channel in Arizona; however, much of thisappears to be due to long-term impacts of discharge fromuranium-mine dewatering operations.
The elevation of the valley floor between the two spurs of
The Hogback is about 6560 ft (2000 m). The 20 ft (6 m)
section of younger valley fill exposed in walls of the Puerco
River "arroyo" from the area of this stop (Fig. 1-2) to central
Gallup (2 mi downstream) comprises three major units: (1)light brown to light yellowish-brown (mostly 10YR hues)
silty clay with thin, fine-sand interbeds and local gravelly
lenses; this unit makes up most of the exposed valley fill;(2) reddish-brown clay and sand that backfill broad channels
cut into and locally bury unit 1; and (3) light brown sandy
alluvium forming the fill of low inset terraces along thepresent Puerco channel.

Leopold and Snyder (1951) briefly studied the stratigraphy of valley-floor alluvium in downtown Gallup and divided the fill into two units. The lower, older alluvium oflate? Pleistocene age is called the Gamerco formation andconsists of reddish and reddish-brown sand with hard calcareous concretions and tubes overlying partially cemented
gravelly sand. A reddish calcareous soil is locally preservedat the top. The upper, younger alluvium is called the Nakaibito formation and consists of sand, sandy silt, and clay
with minor pebbly lenses typical of Holocene alluvium.
Since the local valley fill was first described (in 1946) by
Leopold and Snyder (1951), urban expansion, railroadmaintenance activities, and various river-channel changes
have resulted in obliteration of many features observed in
1946. However, the 20 ft (6 m) "arroyo channel" withdimensions similar to the present one was already present
in 1946, in marked contrast to the 6 ft (2 m) deep channel
observed here by N. H. Darton in 1901 (see photographs in
fig. 2 of Leopold and Snyder, 1951).
The fine-grained valley fill (unit 1) exposed in the 20 ft (6m) arroyo walls between the two spurs of "The Hogback" atthis stop (Fig. 1-3) is part of the Nakaibito "formation" of
Leopold and Snyder. These sediments contain a very sparse
fossil snail fauna that has not been studied. Extensive
systems of "pipes" and related surface-collapse depressions
are present back from the arroyo walls. Underlying older
alluvium (Gamerco formation) is not presently exposed in
the tour-stop area. Overlying channel fills (unit 2) are also

113
and Qoa (Pleistocene). Thick deposits of their older alluvium (Qoa) are seen at Stop 1-3b.
After Stop 1-3a continue northeast 0.45 mi on gradedroad through 1-40 underpass. 9.0 mi, cross North Frontage
Rd. and proceed northeast 0.2 mi on Superman Canyon
Rd. to entrance of a large borrow pit at north edge of
Puerco River valley. Turn right into western part of pit(watch for muddy conditions) and park near base of bluff
for Stop 1-3b.

poorly exposed in this reach of the Puerco. The very young(modern) inset terrace fill (unit 3) is very well expressed
here, however.
Leopold and Snyder (1951) note that the inner valley fillis at least 95 ft (29 in) thick downstream from Gallup, butthey infer that the fill is only about 20 ft (6 m) thick in thisarea. However, no systematic study of valley-fill thickness
or facies distribution has been done in the upper Puerco
River basin (e.g. drilling and geophysical studies), and maximum thickness of valley fill in and upstream from The
Hogback narrows may be much greater than 20 ft (6 m).
In some drainages, Leopold and Snyder (1951) recognized two depositional episodes in the Nakaibito formation.
Pueblo I and II (A.D. 700-1150) potsherds were foundburied 14 ft (4.3 m) below the surface in the type section (inthe Chaco drainage basin to the north!). Few, if any,
radiocarbon dates have been obtained in the Puerco Riverbasin. This stratigraphy is typical in many drainages in theregion, but the names of the formations have not been applied elsewhere even in this drainage basin. Roughly equivalent stratigraphic units (Jeddito = Gamerco, Tsegi =Nakaibito) introduced by Hack (1942) in the Black Mesa
area of Arizona have been mapped in parts of the San Juanbasin (e.g. Blagbrough, 1965; Weide et al., 1979).
Later workers have divided Quaternary alluvium in the
Puerco, Rio San Jose, and Chaco drainages into four mappable units (e.g. Robertson, 1974). The oldest unit, Qoa,
consists of "grayish orange to yellowish gray, poorly topartly consolidated silt, sand, and gravel . . . generally morethan 20 ft (6.1 m) above the level of modern valley flood
plains." On maps within the drainage basin, Qoa occurs at as
many as seven discrete levels ranging up to 1224 ft (373 m)
above the valley floors (Love and Gillam, in press). The next
oldest unit, Qa3, consists of "grayish orange to pinksemiconsolidated silt, sand, and gravel in terraces 5 to 20 ft(1.5-6.1 m) above graded valley flood plains and underlies . .
. surfaces and upper reaches of stream valleys. Generally
underlain by older alluvium." Qa2 is Holocene sand, gravel,silt, and clay of graded valley bottoms and floodplains prior
to latest cycle of entrenchment. It grades laterally tocolluvium. Qal consists of sand and gravel in arroyos, fans,and valley floors that have been deposited since the latest
cycle of erosion (presumably post-1850). In the area of thisstop, Green and Jackson (1976) simply divide the valley-fillalluvium into two broad units, Qal (Holocene)

Stop 1-3b: Older alluvial deposits near South Fork andmainstem Puerco River confluence (J. W. Hawley and
D. W. Love). This stop (elevation range 6580-6660 ft,
2005-2030 m) is located on a large remnant of older
valley-fill alluvium (Qoa map unit of Green and Jackson,1976) overlooking the valley floor at the confluence of themain-stem and South Fork of the Puerco River (0.5 mi, 0.8
km to the east-southeast; elevation about 6560 ft, 2000 m).
Major landmarks of the area are shown in Fig. 1-4. These
include the White Cliffs (2 mi, 3.2 km, to the north), theGallup watergap through The Hogback (1 mi, 1.6 km, tothe southwest), the valley of the South Fork and Fort Wingate Army Depot to the east-southeast, the Red Rocks
cliffs and Red Rock State Park (3.5 mi, 5.6 km, to the east),
Pyramid Rock (3.5 mi, 5.6 kin, to the east-northeast; elevation 7487 ft, 2289 m), and the Church Rock uranium
mine and mill site (12 mi, 19.3 km, to the northeast) up thePuerco River valley. The Entrada Sandstone is probably
the main bedrock substrate at this site (Green and Jackson,
1976; Hackman and Olson, 1977).
As much as 80 ft (24.4 m) of older valley-fill deposits are
exposed in borrow pits at this stop (Fig. 1-2). 5YR hues are
dominant in these deposits in contrast to the 10YR hues inthe younger valley fill seen at Stop 1-3a. An axial-fluvial(ancestral Puerco River) facies is exposed in most of thepits between elevations of about 6580 and 6650 ft (20052027 m). This unit is primarily clean pink to pinkish-white114
sand with thin interbedded layers of light reddish-brownsilt-clay and lenses of sandy gravel. In the highest andnorthernmost pits at this stop (elevations up to 6660 ft,
2030 m) the fluvial facies is overlain by channel gravel(pebbles, small cobbles, and a few boulder-size mud clasts)
representing ancient tributary (arroyo) deposits derived
from uplands in the White Cliffs-Hogback area to thenorthwest. One large fragment of a proboscidean tusk wasnoted in the uppermost part of the fluvial facies.
Original constructional surfaces of the "older" axial-fluvial or tributary alluvial deposits are not preserved here andthe upland erosion surfaces on these units are mantled witha complex of light reddish-brown to brown local alluvial,
colluvial, and eolian deposits. Strongest soils (surficial andshallowly buried) developed in this veneer and immediately
underlying beds of the "older alluvium" have weakly tomoderately developed argillic horizons and weak stage-III
carbonate horizons (as defined by Gile et al., 1966, 1981).
Middle to late Quaternary alluvial chronologies worked
out in adjacent part of the Chaco River basin to the north(see Wells et al. , 1983) and Rio Grande basin to the east(see Days 2 and 3, and supplemental articles, this volume)
support the inference that the thick valley-fill remnant preserved at this stop represents a major, climatically controlled
episode of valley entrenchment followed by an interval ofequally widespread aggradation of at least the inner valley
areas. The youngest possible age for these deposits would befrom early to middle Wisconsin, and could range back to late
middle Pleistocene (including the last pre-Wisconsin fullglacial interval; marine 0-isotope stage 6).
The final major period of widespread valley entrenchment
presumably occurred in the late Wisconsin prior to emplacement of the younger valley-fill units observed at Stop 1-3a.
The time(s) of maximum incision of the deepest parts of the"thalweg" below the Gamerco (Jeddito)-type units in this area
has (have) still not been determined. It is possible that the
deepest entrenchment occurred prior to the deposition of the"older" axial Puerco deposits seen at this stop.
Retrace route to US-66 through ATSF Railroad
underpass south of Stop 1-3a.

upstream from the confluence of the Catalpa Canyon andPuerco River arroyos and the area north of El Rancho Hoteldescribed by Leopold and Snyder (1951) in their discussionof the "type" Gamerco and Nakaibito formations (see Stop1-3a). South of US-66 Boardman Rd. follows Catalpa Canyon, climbing out of the Puerco River valley acrossoutcrops of the Crevasse Canyon Formation (Dilco Coal
Member, Mesa Verde Group). Reclamation work at former
Carbon Coal No. 2 strip mine is visible about 1.5 mi up
Catalpa Canyon south of Boardman Rd. at 3.2 mi.
4.6 mi, traffic light at Boardman Rd.-NM-602 intersection. Turn left and proceed south toward Zuni Pueblo on
NM-602. 5.3 mi, MP 28, roadcuts for next 7 mi in sandstones, shales, and thin coals in Crevasse Canyon Formation. Route crosses irregular, hilly, erosional topography
developed on nonmarine Upper Cretaceous rocks, which
contrasts markedly with the even-crested mesa and cuesta
surfaces formed on extensive marine sandstone units suchas the Gallup and Dakota.
8.3 mi, MP 25, highway skirts arroyo of Bread Springs
Wash to right. Incision of the present 20-25 ft (6-8 m) deepchannel probably occurred during the past two centuries.
9.1 mi, bridge over Bread Springs Wash.
13.3 mi, MP 20, for next 11.5 mi NM-602 crosses a
broad upland erosion surface on Upper Cretaceous rocks;
the surface slopes gently to the west and is cut by shallow
valleys of Puerco and Zuni River tributaries that head on
the western slope of the Zuni Mountains. Locally thickalluvial and eolian deposits of the upper Miocene-lower
Pliocene Bidahochi Formation cap this "Zuni" erosion surface (see Stop 1-4). The unit is mainly poorly cemented
sandstone and mudstone, with local lenses of volcanic ash
and thin, well-indurated beds of siliceous sandstone. 14.3
mi, MP 19, prepare for right turn on Skeets Rd. in 0.3 mi.
14.6 mi, turn right on Skeets Rd., cross cattle guard andproceed west for 2.3 mi to Stop 1-4. Note typical development of piñon-juniper forest. 16.9 mi, park on right sideof Skeets Rd. on crest of interfluve and carefully walk
across road to left to examine Bidahochi exposures invalley to south.

Directions and comments en-route to Stop 1-4
0.0 mi, stop sign at US-66. Turn right and continue west
on Business 1-40. 0.6 mi, traffic light at Patton Rd.
Prepare for left turn at next light (0.7 mi) at Boardman Rd.
1.3 mi, turn left and proceed south on Boardman Rd. (NM564) across Aztec Ave. (traffic light at 1.4 mi). Boardman
Rd. curves to right past Gallup High School (2.2 mi), andcontinues west past Gallup Branch of the University of
New Mexico (to left of four-way stop at 3.9 mi) to junction
with NM-602 at south edge of Gallup. 4.6 mi, traffic lightat intersection of Boardman Rd. and S. 2nd St. Turn left
and proceed south 10 mi on NM-602 to Skeets Rd. 14.6
mi, turn right on graded Skeets Rd. and proceed 2.3 miwest to Stop 1-4.
The US-66 route through eastern Gallup is on the floorof the Puerco River valley and passes the two ridges of The
Hogback through a 0.15 mi (0.25 km) wide watergap. The
route from here to Zuni Pueblo is in the Gallup sag, thenarrow structural depression between the western Zuni andsouthern Defiance uplifts that connects the San Juan andZuni basins.
The US-66-Boardman Rd. intersection (1.3 mi) is justStop 1-4: Bidahochi Formation (D. W. Love and 0. J.
Anderson; modified from Love, 1989b, and Anderson et al.,1989). At this stop about 100 ft (30 m) of Bidahochi
Formation rest unconformably on the Crevasse Canyon Formation. To the west, the Bidahochi Formation is up to 787 ft(240 m) thick and covers over 6178 mil (16,000 km2) inwestern New Mexico and eastern Arizona. The depositional"basin" consisted of paleovalleys and interfluves sloping
southwest and west from Black Mesa, Defiance and Zuni
uplifts, and north from the White Mountains toward theancestral Little Colorado River. After initial erosion of valleys through tilted older rocks, deposition occurred in many
facies, such as alluvial, lacustrine, eolian, monchiquitic
eruptive, travertine and other spring deposits, and silicic ash.
Here on the flanks of the Zuni uplift, only the alluvial,
eolian, and silicic ash facies are apparent (see 19 in. thickash just below road level at this stop). Age relationships witholder and younger Tertiary units are not well constrained.
Parts of the mid-Miocene Fence Lake Formation (McLellan
et al., 1982) may be lateral equivalents or precursors of the
Bidahochi Formation. Deposition near Sanders, Arizona,started by 9 Ma and perhaps as early as 12 Ma based on
fossil evidence (Love, 1989b). Aggradation

11<of Bidahochi Formation ceased by about 4 Ma in mostareas. Pliocene and Pleistocene terraces are inset into the
Bidahochi Formation. In general, depositional setting,facies, and age of the Bidahochi are similar to the Ogallala
Formation in eastern New Mexico and northwestern Texas,
implying similar tectonic, geographic, and climatic reasonsfor both accumulations.
In this area, McCann (1938) called the erosion surface atthe base of the Bidahochi the "Zuni" surface. To the northand west, other erosion surfaces beneath other units orwithin the Bidahochi Formation have been correlated withthe "Zuni" surface (cf. Akers, 1964; Cooley and Akers,
1961). Much of the correlation of erosion surfaces was done
before the advent of radiometric ages. Much work remains
to be done to revise the nomenclature of unconformable anddisconformable Tertiary surfaces in the Little Colorado
drainage basin.
Retrace route to NM-602.
Directions and comments en-route to Stop 1-5
0.0 mi, stop sign on Skeets Rd. at NM-602. Turn rightand proceed south for 14.2 mi on NM-602 toward Zuni
Pueblo through Vanderwagen (Vander Wagen of Pearce,
1965; 6.3 mi). Bear right on Zuni Hwy 4 half a mile southof MP 5 and continue south for 4.7 mi to NM-53. Turnright on NM-53 and continue west for 6 mi through Black
Rock (20.7 mi) to central Zuni Pueblo (25 mi) and accessto Stop 1.5.
0.0 mi, for the next 10 mi the route continues south on
NM-602 across broad upland surfaces capped with Bidahochi Formation. Cretaceous bedrock is only exposed in a
few valleys of major streams. 0.5 mi, Bread Springs-Pine
Haven Rd. to left is route of New Mexico GeologicalSociety Field Conference tour (Anderson, Love, and Lucas,
1989) to McGaffey-Fort Wingate area (Stops 1-1 and 1-2).
Crossing shallow upper valley of Manuelito Wash (Puerco
tributary). 0.7 mi, MP 18, Joe Milo's Trading Post to right.
Elvis reportedly appeared here in September 1986 (Anderson et al., 1989). 1.7 mi, MP 17, white silicic ash bed in
Bidahochi exposed in roadcut ahead to right. 3.7 mi, MP
15, road to right leads to BIA elementary schools at Chichiltah and Jones Ranch. Route ahead crosses shallow
upper valley of Whitewater Arroyo (5.1 mi, Puerco
tributary), with local outcrops of Crevasse Canyon
Formation/Gallup Sandstone "basement complex." 6.3 mi,
Vanderwagen and Whitewater Trading Post to right. Valleyof Nelson Wash ahead with exposures of Gallup Sandstone.
6.7 mi, MP 12, drainage divide between Puerco andZuni river basins capped with Bidahochi Formation. 8.0
mi, McKinley County Rd. 6 to right to Kiwanis Park andCousins Brothers Trading Post. 9.9 mi, McKinley County
Rd. 8 to right. Entering Zuni Indian Reservation.
10.2 mi, leaving upland surface capped with Bidahochi
Formation. View to south down valley of tributary to Rio
Nutria, which joins Rio Pescado to form Zuni River about 8
mi (13 km) downstream. As highway drops into the valley,note dip-slope to right formed on Dakota Formation. Strikevalley ahead is eroded into Rio Salado (Shale) Tongue of
Mancos Formation; and escarpment to left capped by Atarque (Sandstone) Member of Tres Hermanos Formation.
Dips here are 3-5° to the east-northeast.
11.7 mi, MP 7, watershed management practices implemented by local U.S. Soil Conservation Service staff have

led to successful reclamation of arroyos in this segment of
the Rio Nutria drainage system. Prominent cliffs forming
eastern valley wall are Atarque (Sandstone) Member overlain by carbonaceous Carthage Member of Tres Hermanos
Formation. Twowells (Sandstone) Tongue of the Dakota
overlies Whitewater Arroyo (Shale) Tongue of the Mancos
in a small hill at left edge of highway. 13.7 mi, MP 5, keepright for Zuni Hwy 4 in 0.5 mi.
14.2 mi, bear right on Zuni Hwy 4 and proceed south tovalley of Zuni River across cuesta dip-slope capped withsandstones of the Dakota Formation. Roadcut exposuresahead in carbonaceous shales near base of Dakota. 15.7 mi,crossing west-facing cuesta rim. 15.9 mi, contact of Dakota
on Jurassic Zuni Sandstone (Cow Springs/Entrada Sandstone equivalent with Todilto interval missing). 17.2 mi,crossing thick Quaternary eolian sand sheet at base of escarpment. Broad alluvial toeslope ahead, with eolian
veneer, is graded to floor of Zuni River valley, which here
coincides with surface of Pleistocene basalt flow east ofBlack Rock (about 3 mi, 5 km, to the southwest). 18.6 mi,bridge over Zuni River, which is shallowly entrenched into
basalt flow of Rio Pescado lobe (North Plains section of
the ZuniBandera field). The tip of this flow forms Black
Rock promontory 2 mi to the west. A K-Ar age determined
here by Laughlin et al. (1982) is 0.7 ± 0.55 Ma (olivine
tholeiite sampled about 0.7 mi upstream). Prepare for rightturn ahead.
18.9 mi, stop sign at NM-53. Turn right on NM-53 andproceed west toward Black Rock and Zuni Pueblo. 19.7 mi,
MP 17, prominent mesa from 11:00 to 12:00 is Dowa Yalanne(Corn Mountain) capped by Dakota Sandstone which overlies
about 500 ft (150 m) of Zuni Sandstone. Bosque to right is ondelta of Zuni River at upper end of Black Rock Reservoir. The
dam is about 1 mi to the west-northwest. Kirk Bryan examined
one of its failures in 1928 and made recommendations on dam
reinforcement. He summarized the Quaternary history of thearea in his unpublished report. 20.7 mi, MP 16, Black Rock
community to right on basalt flow. Chinle redbed exposuresahead on right and left.
21.9 mi, basalt of Black Rock caps Chinle Formation
and/or Chinle-derived alluvium to right at end of Rio Pescado lobe of North Plains flow. Continue west on NM-53
across low scarp to floor of broad alluvial valley underlain
by Chinle Formation. Orr (1982) recorded as much as 236
ft (72 m) of alluvial fill in the valley to the west. 22.7 mi,
MP 14, Zuni Pueblo Historical Marker on right. Corn
Mountain road to left (22.8 mi). See Chamberlin et al.
(1989) and Lucas and Anderson (1989) for overview ofZuni area geology and history.
23.2 mi, NM-53 bridge over Zuni River. 24.0 mi, caution!
School Crossing; Zuni High School to left. 24.6 mi, congested area; Tribal Headquarters and shopping center toright; older part of Pueblo to left. 25.0 mi, four-way stop.
Access to west and north to gullied valleys of the Zuni
area, including Bosson Wash (Optional Stop 1-5). Please
note that all visits to Zuni tribal lands outside the central
Pueblo area must be done with permission from the Tribal
Governor's Office and other appropriate officials.
Stop 1-5 (optional): Arroyo processes and valley-fill sequences: The impacts of Holocene climatic changes,
and land and water management practices. Comments
and discussion by S. G. Wells. Retrace route on NM-53 tojunction with Zuni Hwy 4 (0.8 mi east of MP 17).

116
Directions and comments en-route to Stop 1-6
Route is eastbound on NM-53 from junction with Zuni
Hwy 4. Continue 41.2 mi east on NM-53, via Ramah andEl Morro National Monument to junction with U.S. Forest
Service Road 50 (FS-50; 0.1 mi west of MP 59). Turn left
and proceed 1.5 mi north on FS-50 to junction with FS187. Turn left on FS-187 and continue about 1.4 mi northwest to Stop 1-6 at Oso Ridge Lookout Tower on summitof Oso Crater. Four-wheel drive is usually needed to reachthe summit, so park other vehicles on straight section of
FS-187 below switchback about 0.5 mi from lookout tower
and walk to Stop 1-6.
0.0 mi, Zuni Route 4 to left; continue east on NM-53 up
valley of Zuni River through watergap in east-dipping cuesta
developed on Dakota/Zuni sandstones. 0.3 mi, MP 18, radio
towers on skyline at 10:00 are in area of cuesta summit
where patches of siliceous gravels on the Dakota Sandstone
caprock have been noted. Age and origin of these andsimilar very high-level deposits elsewhere in the Zuni
Mountain region have not yet been determined. However,they could be as old as Eocene and as young as Pliocene. 1.3
mi, MP 19, Rio Pescado tongue of North Plains basalt flow
to left along river. Spectacularly banded exposures of Zuni
Sandstone in cliff faces noted by Dutton (1885).
2.3 mi, MP 20, head of Zuni River to left at confluenceof Rio Nutria (from the north) and Rio Pescado. Continue
east up valley of Rio Pescado. 3.3, MP 21, major junctions
ahead: NM-602 to left (3.4 mi) and NM-36 to right (3.9mi); continue east on NM-53 toward Ramah and El Morro
National Monument.
4.3 mi, MP 22, cliff face from 9:00 to 11:00, capped by
Gallup Sandstone, is reference section for Tres Hermanos
Formation. Lower part of valley cut in Mancos Shale tongues.
This northeast-dipping marine (intertonguing sandstone andshale) section disappears below road level over the next 3.3
mi. 7.6 mi, nonmarine (upper) Gallup Formation exposed invalley walls and tableland escarpments ahead for next 5 mi.
This coal-bearing part of the section is characterized byreddish, coarse, feldspathic sandstones.
9.9 mi, cross bridge over Rio Pescado. Village ofPescado ahead. 10.8 mi, at 10:00 to 11:00 Rio Pescado
lobe of North Plains basalt flow in stream bottom. 13.0 mi,leave Zuni Reservation.
14.3 mi, MP 32, descend into Ramah Valley of Rio
Pescado. 14.8 mi, Ramah watergap at 10:30 exposes lightcolor-banded cliffs of Zuni Sandstone with darker Dakota
Sandstone capping dip-slope of cuesta-forming southwestern flank of Zuni Mountains. 15.4 mi, enter Ramah, settled
by LDS pioneers in 1874 and named after a figure in theBook of Mormon. 16.3 mi, MP 34, leave Ramah.
16.9 mi, low sandstone bench on valley floor to left is
Dakota Sandstone. 18.0 mi, McKinley-Cibola County line.
Road follows the southwestern edge of the Zuni uplift in a
broad synclinal valley cut in Mancos Shale. Forested mesa
from 1:00 to 4:00 capped by sandstones of Tres Hermanos
Formation.
18.6 mi, tree-covered dip-slope on Dakota Sandstone toleft. 20.3 mi, MP 38, Lewis Trading Post to left. 21.0 mi,
crossing lobe of basalt from Rio Pescado tongue for about
next 2 mi. 22.3 mi, MP 40, pinnacles of Zuni Sandstone inwatergap at 9:00 are "Los Gigantes"; note crest of Zuni
Mountains (Oso Ridge) in distance beyond cuesta cappedwith Dakota/Zuni Sandstones.
23.2 mi, roadcuts in Two Wells Sandstone Tongue of

Dakota. At 11:00, Inscription Rock (El Morro National
Monument) is in Zuni Sandstone cliff at east end of mesa
capped by Dakota Sandstone. Basalt cinder cone of North
Plains section of Zuni-Bandera field in distance at 12:00.
23.9 mi, Navajo Hwy 125 to right leads to Ramah Agency
headquarters. 24.5 mi, for next 8 mi basalt flows of thewestern Zuni-Bandera (North Plains) field cover most of thevalley floor. These flows may be as young as late middle
Pleistocene (Laughlin et al., 1982). 25.7 mi, enter El Morro
National Monument. 26.3 mi, MP 44, Zuni Sandstone cliff
to right. Inscription Rock at base of cliff (26.5 mi) has theoldest known Spanish inscription in New Mexico: Don Juan
Oñate, 16 April 1605. 27.2 mi, road to El Morro National
Monument Visitor Center to right. The tour route is back in
the Acoma-Zuni (physiographic) section from here to the
Laguna Pueblo area east of Grants (Day 2, Stop 2-1).
27.8 mi, roadcut exposures of North Plains basalt flow.
28.2 mi, Inscription Rock RV Park to right. Crest of Zuni
Mountains (Oso Ridge) forms northern skyline. 29.2 mi,more roadcuts in North Plains flow. 30.6 mi, cinder conesin El Malpais and Chain of Craters (North Plains) areas of
the Zuni-Bandera field at 12:30 to 3:00. 31.3 mi, MP 49,
Oso Ridge, on northern skyline, is composed of Glorieta
Sandstone and overlying limestones and minor sandstonesof the San Andres Formation. The Continental Divide follows this segment of Oso Ridge and the chain of cindercones to the east and south.
31.8 mi, Tinaja Bar and Trading Post on right. 33.3 mi,
MP 51, for the next 6 mi the route skirts the northern edgeof the North Plains section of Zuni-Bandera field. The
alkali-basalt flow here is locally buried by upper Quaternary
sediments (alluvial, eolian, lacustrine), which also lap onto
the southward-dipping Permian and Triassic rocks thatflank the southern Zuni uplift. Features of the broad strikevalley (mainly cut in Chinle Formation) that parallels thetour route are completely masked by a cap of PlioceneQuaternary basalts. 33.8 mi, road to left to limestone quarry
in San Andres Formation.
36.3 mi, MP 54, red-maroon beds exposed at 10:00 are
sandstones of Moenkopi and basal Chinle Formations.
37.3 mi, MP 55, lava cone to right. Bandera Crater (left)
and Cerro Bandera (right) cinder cones at 12:00 about 5 mi
(8 km) to east-southeast. 38.8 mi, Oso Ridge cinder cone
(Stop 1-6), on skyline at 10:00, is on the southeast end of
Oso Ridge. 40.3 mi, MP 58, Oso Ridge cone (9:00), Bandera Crater (1:00), and Cerro Bandera (1:30) are all on theContinental Divide. K-Ar age of alkali-basalt flow sampled
along NM-53 about 1 mi northwest of Bandera Crater (nearMP 59) is 0.199 ± 0.042 Ma (Laughlin et al., 1982).
Prepare for left turn in 0.9 mi.
41.2 mi, turn left on Forest Service Road 50; proceed
north on FS-50 up cuesta dip-slope on San Andres
Limestone and Glorieta Sandstone. 42.4 mi, crossing
Continental Divide (elevation 8190 ft, 2496 m). 42.7 mi,turn left on FS-287 to Oso Ridge Lookout Tower on Oso
Ridge cinder cone (8713 ft, 2656 m). Walk to Oso Ridge
lookout tower and Stop 1.6.
Stop 1-6: Southern El Malpais area of Zuni-Bandera
volcanic field area from Oso Ridge cinder cone (A. W.
Laughlin, J. W. Hawley, and D. W. Love). The summit of
the Oso Ridge cinder cone offers a broad vista of the region
south of the broad Zuni Mountain summit area. Mt. Sedgwick (9256 ft, 2818 m) 10 mi (16 km) to the north, and117
the crest of Oso Ridge (>9000 ft, 2743 m) to the northwest
are high points of the uplift (Fig. 1-5). The route to Stop 17 crosses the Precambrian core of the Zuni uplift anddescends to the Rio San Jose valley at Grants through Zuni
Canyon, which heads about 8 mi (13 km) to the northeast
at the prominent gap northwest of San Rafael Mesa (Fig. 15). Basalt from the Paxton Springs cone, 3.5 mi (5.6 km) to
the east-northeast, flowed through this canyon onto theplain southwest of Grants in latest? Pleistocene time.
This tour stop provides an overview of the local volcanic
terrane (western El Malpais and North Plains of the ZuniBandera field; Laughlin et al., 1982) and ongoing work inlocal ice caves on their potential as sites for paleoclimatological research. The Pliocene-Quaternary geology of thenorthern and eastern part of the plains area (El Malpais
lava field of Maxwell, 1982, 1986) and emplacement of thelate Holocene McCartys flow are discussed in more detailat Stop 1-7.
An ancient upland erosion surface, which probably
started to form as early as in the Eocene, is preserved
beneath the extensive flows that mantle the plains to thesouth and east of this stop. Significant Oligocene to
Pliocene extensional tectonic and erosional features may
underlie the North Plains area, particularly east of theContinental Divide (Fig. 1; Chain of Craters belt, Fig. 1-5).
Post-Laramide evolution of the southeastern margin of theColorado Plateau is discussed by Cather (1989) and
Chamberlin and Anderson (1989b).

According to Baldridge, Perry, Vaniman, et al. (1986, p.
196):
The oldest basalt recognized in the Zuni-Bandera field is
3.8 Ma [sampled at southern base of Cerro Rendija, Fig. 15] (. . . Laughlin et al., 1982). Most of the basalts presently
exposed in the . . . field are Quaternary in age. Older Quaternary rocks were erupted mainly from a group of 30 cones
called the Chain of Craters . . . Younger Quaternary basalts
were erupted mainly northeast of the chain, forming an area
referred to [by Maxwell, 1982, 1986] as El Malpais . . . On
geomorphic evidence, Gawell (1975) and Gawell and
Laughlin (1975) recognized a gradual decrease in age fromsouthwest to northeast along the northern portion of the
Chain of Craters. A progressive change from alkalic basalts
to tholeiites . . . accompanied this change. Elsewhere in the
ZuniBandera field, compositions range from basanites totholeiites.

The following comments on the Bandera Crater area by
Baldridge, Perry, Nealy, et al. (1989, pp. 217-218) introduce a discussion of laminated ice bodies in collapsed lava
tubes (ice caves) at El Malpais National Monument by
Thompson et al. (this volume):
Bandera Crater is a breached symmetrical cinder cone about
1 km [0.6+ mi] in diameter and 150 m [500 ft] high. The
central depression is about 180 m [600 ft] deep, extending
80 m [260 ft] below the level of the breach. A large lava
tube, which extends at least 20 km [12.4 mi] to the south, is
exposed in the breach. Although collapsed for the first 800m [2625 ft], most of the remainder of the tube is intact.
After the McCartys flow [Stop 1-7], the flows fromBandera Crater are the youngest in the Zuni-Bandera [El
Malpais] lava field. The [alkalic basalt] lavas flowed southand east around the Zuni Mountains to a point of 12 km[7.5 mi] south of Grants. The Bandera flows overlie 0.199
Ma basalt [near MP-59, NM-56; Laughlin et al., 1982] but
are themselves too young to date by the K-Ar method . . .
The Ice Caves are located in the main lava tube fromBandera Crater. Here, because of the overhang of thecollapsed lava tube and the insulating properties of basalt,
ice that forms in the winter is preserved throughout theyear. The green color of the ice is due to presence of algae.
In frontier days, ice from the caves was transported tonearby Fort Wingate [original site near San Rafael]. Other
ice caves occur in lava tubes of this area, but are not open to
the public.

Return to junction of Forest Service Roads 50 and 187.
Directions and comments en-route to Stop 1-7
This log is adapted primarily from Colpitts et al. (1989).
0.0 mi, proceed north on FS-50 from its junction with FS187. 2.9 mi, FS-50 turns left towards McGaffey (Stop 11) atjunction with FS-49. Continue straight ahead on FS-49
across summit of Zuni Mountains and through Zuni Canyon
to southwest edge of Grants. 14.9 mi, cattle guard; leaving
Cibola National Forest. Continue down paved Zuni Canyon
Rd. 18.4 mi, Grants city limits. 19.2 mi, sharp right turnfollowed by sharp left turn, passing Zuni Canyon Trading
Post on left. 19.3 mi, stop sign at end of Zuni Canyon Rd.;rejoin NM-53, turning left towards Grants. Prepare forright turn ahead onto 1-40. 19.5 mi. Take eastboundonramp to 1-40. 20.6 mi, MP 83, proceed 6.2 mi east on
1-40 to Exit 89. 26.8 mi, take Exit 89. 26.9 mi, stop sign at
NM-117; turn left, cross over 1-40, and park at westedge of Stuckeys Service area for Stop 1-7.

0.0 mi, proceed northwest on FS-50, descending cuesta
escarpment cut on Glorieta Sandstone. Cross concealed Glorieta-Yeso contact beneath hillslope deposits at switchback(0.2 mi). At 12:00 cinders from Oso Ridge cone rest on

118
upper Yeso beds. 0.5 mi, base of basalt flow from Oso
Ridge vent rests on cinder deposit to left. Flow continues
down valley ahead (11:30 to 12:00).
1.0 mi, "No Easy Life Ranch" entrance on left. Low ridge
to right (2:00 at 1.1 mi) is formed by Meseta Blanca
Sandstone Member of Yeso Formation, which also underlies
the Oso Ridge flow to left for the next mile. Road is primarily on thin cover of upper Quaternary valley fill with afew exposures of Meseta Blanca Sandstone. Entering area of
detailed mapping by Goddard (1966).
2.0 mi, on Oso Ridge flow with discontinuous sedimentary cover for next 0.3 mi. Quaternary units conceal Yeso/
Abo redbed contact. Paxton Springs road to right (2.1 mi).
Cattle guard and Cibola National Forest boundary (2.2 mi).
2.3 mi, Abo Formation exposed in roadcuts at north edge of
Oso Ridge flow; route on Abo redbeds for next 0.7 mi.
2.7 mi, cattle guard. At 2:30, Paxton Springs cinder cone
and associated basalt flows (latest? Pleistocene). This is thesource of the flow in Zuni Canyon and southwestern Grants
area (9.2-21.6 mi). 2.8 mi, crossing old railroad grade, partof an old logging route extending through Zuni Canyon. 2.9
mi, junction; continue straight ahead on FS-49. Arkosic
sandstones and pebble conglomerates near base of Abo Formation are exposed in roadcut ahead. 3.0 mi, bridge over
Agua Fria Creek. Contact of Abo Formation on deeply
weathered and fractured granitic gneiss (Precambrian) along
creek bed about 200 yards upstream (Stop 1 of Colpitts etal., 1989). Cinder cones marking Paxton Springs vent 1 mito the east-northeast (2:30). On upper Quaternary valley fill
for next 0.2 mi.
3.2 mi, cross exhumed erosion surface cut on granitic
gneiss forming core of Zuni uplift for next 3 mi. Basalt
from the Paxton Springs vents caps this surface to the eastof FS-49. 5.0 mi, road ascends to surface of small western
lobe of Paxton Springs flow. Note pressure ridge to right.
For the next mile, route is mainly on valley fill andparallels old logging railroad grade to right.
6.4 mi, back on granitic gneiss; approaching Abo/Precambrian contact in road ahead. 6.5 mi, possible marine
limestone of Pennsylvanian age (pre-Abo?) is exposed inridge to left (Colpitts et al., 1989, Stop 2). 6.7 mi cattle
guard. Basal Abo sandstone and conglomerate to left. 7.0
mi, at 11:00 fault contact between Abo and Precambrian,with head of Limekiln Canyon forming break inescarpment on skyline.
7.2 mi, cross bridge in lower La Jara Canyon; on valley
fill for next 1.8 mi. Abo redbeds form low ridge in middle
distance. High ridge on skyline is escarpment of cuesta
capped by northeast-dipping San Andres Formation (limestone, sandstone, and dolomite). The San Andres is underlain by light-colored, quartzose sandstones of the Glorieta
Formation, which in turn overlies pale reddish-brown toreddish-gray sandstones and dark-gray limestones of theYeso Formation. 7.8 mi, Paxton Springs flows on valley
floor at head of Zuni Canyon from 2:00 to 3:00. Cuesta
summit to east of canyon is called San Rafael Mesa.
8.3 mi, junction with Forest Service Road 480 (to left).
Continue straight ahead on FS-49. Basalt flow ahead andto left is a "pre-Paxton Springs" unit derived from a venton the west side of Cerro Colorado (peak of Precambriangneissic aplite 2 mi to the northwest). Lobes of this flow("Zuni flow" of Maxwell, 1986) also extend down canyons
in the cuesta escarpment north and east of Cerro Colorado(e.g. Limekiln Canyon). 8.5 mi, steeply dipping Abo strata

occur adjacent to fault separating Abo and Yeso Formations.
8.7 mi, end of "Zuni flow" to left. 8.8 mi, Stop 3 (upper Yeso
Formation stratigraphy) of Colpitts et al. (1989).
9.2 mi, junction with Forest Service Road 447 (to right)
at head of Zuni Canyon. Proceed down canyon on FS-49,
with remnant of Paxton Spring flow to left. Permian units
exposed in 600 ft (190 m) canyon walls ahead are (ascending) Yeso, Glorieta, and San Andres Formations. 9.8 mi,old logging slide at 3:00. Logs cut on San Rafael Mesa
were run down the slide by cable and used as railroad ties.
10.4 mi, at top of cliff at 9:00 note exhumed pre-Triassic
karst features in upper part of San Andres. 10.6 mi, remnants of Paxton Springs basalt flow on both sides of road.
12.4 mi, canyon crosses broad anticlinal flexure. 12.7 mi,cattle guard. 13.0 mi, basalt crops out in canyon bottom toleft. Yeso redbeds at 3:00 in canyon wall. At 10:00, threequarters of the way upslope, is the San Andres/Glorieta
contact. Basalt flow crops out in canyon bottom at 9:00.
13.7 mi, basalt in roadcut to left.
14.9 mi, cattle guard; leaving Cibola National Forest;
continue on paved Zuni Canyon Rd. Note basalt rims alongbase of canyon walls ahead. 15.2 mi, canyon now only in
San Andres/Glorieta Formations. 15.7 mi, last canyon exposures of Paxton Springs flow ahead. Flow is buried bylate Quaternary (Holocene?) valley-fill deposits down canyon. 16.7 mi, San Andres-Glorieta contact in lower canyon
walls. 16.9 mi, Mount Taylor on skyline at 12:00 is flanked
to the north by La Jara Mesa and to the south by Horace
Mesa. 17.4 mi, cattle guard. Limestone quarry in San
Andres Formation ahead on left; approaching mouth ofZuni Canyon.
17.9 mi, on alluvial plain at canyon mouth; valley fillthins to east and large areas of aa-lava surface on Paxton
Springs flow are exposed. 18.4 mi, Grants city limits. 19.2
mi, sharp right followed by sharp left turn. 19.3 mi, stop
sign at NM-53. Turn left (north) on NM-53; keep to rightand take onramp to 1-40 East (19.5 mi).
20.0 mi, eastbound on 1-40. Spring mound at 2:00, surrounded by distal part of Paxton Springs flow. 21.6 mi, MP
84, back on older Pleistocene flow from El Calderon center.
22.4 mi, Exit 85 to motel area in east Grants. Continuesoutheast on 1-40 across northern edge of El Calderon flow
and adjacent alluvial flats of the Rio San Jose. 23.4 mi,Rancho del Padre Spring is in small draw to right, and othersprings and seeps are in the marsh area to the left (White,1989). Spring discharge between here and the Acoma Reservation boundary (about 5 mi, 8 km, downstream) contributes much of the base flow and good-quality water to the
Rio San Jose.
23.6 mi, MP 86, a faulted Pliocene basalt flow caps
Horace Mesa at 9:00 to 11:00 (see Drake et al., thisvolume). Mesa escarpments are nearly continuously
mantled with landslides and colluvium. East-dipping Upper
Cretaceous rocks of the Dakota-Mancos-Gallup-Crevasse
Canyon sequence are poorly exposed in most areas. 26.6
mi, MP 89, highway cuts across a narrow "neck" of upper
Holocene McCartys flow. Keep right for next exit. 26.8 mi,take Exit 89. Note basal Dakota Sandstone and small thrustfault in roadcut to right. 26.9 mi, stop sign. Turn left on
NM-117 for Stop 1-7 at Stuckeys Service area northwest of
1-40 Interchange.
Stop 1-7: Northern El Malpais area of the Zuni-Bandera
field from the McCartys flow (D. W. Love, J. W. Hawley,

119
and P. G. Drake). The McCartys basalt flow at this stop
overlies the steep, east-dipping Grants monocline (western
limb of the McCartys syncline) within the Acoma sag eastof the Zuni uplift (Figs. 1-6, 1-7). Dakota Sandstone caps
Las Ventanas Ridge to the south and the low ridge at El
Canoncito of the Rio San Jose (0.5 mi to the northwest).
The broad, lava-covered plains to the southwest between
the east end of the Zuni Mountains and Las Ventanas
RidgeCebolleta Mesa (Fig. 1), have been inundated bybasalts of the El Malpais and North Plains (Zuni—Bandera
field) numerous times during the Pliocene and Quaternary.
The most recent event was emplacement of the Upper
Holocene McCartys flow. The flow, originally describedby Nichols (1946), is about 30 mi (48 km) long andoriginated at a small cinder cone about 23 mi (37 km)
south-southwest of this stop (Fig. 1-6; Maxwell, 1986).
The flow is about 5 mi (8 km) wide at its source, but isconfined in a narrow (as little as 60 m) valley immediately
southwest of this stop. Geomorphic expression andstratigraphic relationships with respect to adjacent valleyfill deposits (alluvial, lacustrine, and eolian) indicate thatthe flow is about 1000 years old. Thus the after A.D. 700
age estimate of Nichols (1946), based on Indian legends
and indirect archaeological evidence, is quite reasonable.
There is now no integrated surface runoff, and no namehas been given to the former major tributary of the Rio San
Jose that cut this valley buried by the El Malpais flows.
One rancher at the southern (upper) end of the plains re-ported drilling through about 900 ft (275 m) of basalt flows
before reaching gravel overlying bedrock (Triassic?; C.
Maxwell, pers. comm. 1988). Kelly and Reynolds (1989)
show about 600 ft (183 m) of fill at the lower end of thevalley about 4 mi southwest of this stop. Their seismicprofile interpretation (Fig. 1-8) indicates that the basalts,
ancestral Rio San Jose fluvial deposits, and other valley-fillfacies (eolian, lacustrine) occupy a moderately deep structural basin (graben) of late Cenozoic age.
Flows exposed on the valley floor have K—Ar ages
ranging back to 3.8 Ma (Laughlin et al., 1982; Stop 1-6). The
older reported ages are questionable at least for flows in thenortheastern part of the field, because flows could not havebeen emplaced near the present valley-surface level before
extrusion of the high-level Horace Mesa flows 2.5 to 2.9 Ma(Lipman and Mehnert, 1980) and possibly correlative basalts
capping Mesa Negra (Fig. 1-6) and Cebolleta Mesa south of
this stop (Maxwell, 1986). Two younger flows with

K—Ar ages of about 1.5 Ma (Laughlin et al., 1982) may
indicate that maximum valley entrenchment occurred prior
to that time. Maxwell (1986) and Kelly and Reynolds(1989) show a young fault along the southeastern margin of
the valley (Fig. 1-8); but even if subsidence has occurred
during relatively recent tectonic movement, it is unlikelythat basalts as old as 3.8 Ma could have been displaced
more than 1000 ft (300 m) to form the present valley.
Valley margins along the El Malpais—North Plain flows
have not eroded back appreciably since the flows' emplacement. Numerous landslides, talus cones, and alluvial—colluvial aprons grade to the present valley base level. The lackof extensive hillslope retreat suggests that mass-wastingprocesses also take place episodically and, on the average,
rather slowly, over long intervals. See Drake et al. (thisvolume) for additional information on the area of this stop.
Return to Grants via 1-40; take Exit 85 to motel area.
End of Day 1.

Day 2: Grants to Estancia basin
Summarydrainage system in the Colorado Plateau—Basin and Range
Part 1 of today's tour traverses the southeastern Colorado transition zone is discussed at Stop 2-1 near Laguna Pueblo
Plateau (Acoma—Zuni section, Fig. 1) and the central Rio (Fig. 1). Stop 2-2 introduces the structural framework of theGrande rift area of the Basin and Range province (Mexican Albuquerque basin (central Rio Grande rift), and volcanic
Highland section, Fig. 2). Emphasis is on the Pliocene— units and basin-fill (Santa Fe Group) stratigraphy in the
Quaternary geomorphic evolution of the lower Rio San Jose south-central part of the basin west of Los Lunas. Emphasis
and Rio Puerco valleys, and the southern Albuquerque basin. at Stop 2-3 is on late Quaternary geomorphic processes andvalley-fill stratigraphy in the Rio Puerco valley
Volcano-tectonic and geomorphic evolution of San Jose

121
west of Belen. Optional Stop 2-4 on the Rio Grande valley
rim overlooking Belen focuses on the hydrogeologic framework of the upper Santa Fe Group (Plio-Pleistocene) aquifer
system. Discussion at Stop 2-5 at the Sevilleta National
Wildlife Refuge in southeastern Albuquerque basin emphasizes late Holocene ecologic changes at the northern end of
the Chihuahuan Desert. This locality is also near the boundary between the Basin and Range (Rio Grande rift) andsouthern Great Plains structural provinces.
Part 2 of today's tour (as well as the first stop of Day 3)
is in the closed Estancia basin east of Mountainair (Fig. 2),

which is the site of late Pleistocene Lake Estancia (Meinzer,
1911). Tour stops (E-1, E-2, E-3) emphasize lake and playa
features of late Wisconsin and Holocene age. Estancia basinand highlands to the south and east also mark a physiographic transition from the Basin and Range province (Sacramento section) to the southern Great Plains (Pecos valley
section).
Road log entries between Grants and Mountainair are
primarily compiled from Hawley, Crumpler, and Wells (1982);
Hawley, Love, and Wells (1982); and Maxwell et al. (1989).

Day 2, Part 1: Grants to Mountainair
Directions and comments en-route to Stop 2-1
From the parking lot of the Inn at Grants go south (0.5mi) to 1-40 Interchange (85) and east on 1-40. 1.7 mi, MP
86, crossing the El Calderon flow. 4.9 mi, Exit 89 to right(Day 1, Stop 2-7), crossing narrow neck of McCartys flow
confined in a former arroyo by Las Ventanas Ridge (east,
Morrison and Dakota Formations) and Pleistocene El Calderon flow (west). Route crosses El Calderon and McCartys
flows for next 4.8 mi. 6.7 mi, MP 91, crossing Rio San Jose;for next 2 mi route crosses distal segment of McCartys flow.
Note pressure ridges and water-filled depressions ofcollapsed lava tubes. 8.7 mi, MP 93, rest area ahead (0.2 mi)on McCartys flow. 9.7 mi, MP 94, end of McCartys flow;route ahead on upper Holocene fill of Rio San Jose valley.
Bridge over Rio San Jose; exposures before railroad
underpass are Whitewater Arroyo Tongue of Mancos Shaleoverlain by Twowells Tongue of Dakota Formation.
12.7 mi, MP 96, a small, steep dome is visible in middle
distance in front of Mt. Taylor at 10:00-10:30 where beds ofCrevasse Canyon Formation dip steeply away on all sides.
A drill hole in the center of the dome encountered a mafic
intrusive just below the surface. Broad cuestas to right arecapped by Cubero Sandstone Tongue of the Dakota over the
Oak Canyon Member; a veneer of high-level terrace gravelis commonly present. 15.7 mi, MP 100, Acomita Dam andreservoir to right; at 1:00 is Flower Mountain volcanic center. 17.0 mi, broad, flat area adjacent to 1-40 on right isinterpreted to be the site of a small Pleistocene lake in an
erosional depression in the Mancos Shale. 18.7 mi, MP
103, ahead, route descends through Cubero and Oak
Canyon Members of Dakota Formation.
20.2 mi, take Exit 104 to Cubero, Budville, and Paraje.
Turn left across 1-40, then turn right (east) toward Paraje on
NM-124 (mi 20.6). Dakota/Morrison contact exposed near topof mesas on both sides of valley. Numerous landslides
covering Morrison Formation below. 23.3 mi, downtown
Paraje, road to right leads to Casa Blanca and Acoma Pueblo.
24.9 mi, turn off to Laguna—Acoma High School on left. A
well near here pentrated 160 ft of alluvium. Stratigraphy of
Frog Mesa (9:30-11:00) from top down: basalt flows from
Wheat Mountain cone overlying Clay Mesa Tongue of
Mancos Shale, light-colored Cubero Tongue of Dakota
Sandstone, paludal shale of Oak Canyon Member of Dakota,
and dark-gray sandstone ledge of Encinal Canyon Member of
Dakota. Landslide debris covers the Brushy Basin Member ofthe Morrison Formation. The surface be

low the basalt flows is called the Wheat Mountain surface
by Moench and Schlee (1967) and represents an early stageof valley entrenchment by the Rio San Jose. The youngest
flow on Frog Mesa has a K—Ar age of 2.42 Ma (Lipman
and Mehnert, 1980). 25.1 mi, buried well-developed soil inarroyo cuts on both sides of bridge. 26.2 mi, MP 22 and
New Laguna Post Office. Landslides over eolian sandstoneof Morrison ahead. 27.9 mi, NM-279 to left goes toJackpile uranium mine and villages of Paguate, Bibo, andSeboyeta. 28.1 mi, cross AT&SF Railroad. Prepare to turnleft for Stop 1. 28.3 mi, turn left toward old baseball field,park, and walk to railroad cut.
Stop 2-1: Geomorphic and tectonic evolution of thelower Rio San Jose drainage basin. See Drake et al. (thisvolume) for tour-stop discussion.
Directions and comments en-route to Stop 2-2
Entries for this part of the trip are modified from Maxwell
et al. (1989), Hawley, Love, and Wells (1982), Hawley,
Crumpler, and Wells (1982), Wells et al. (1983), and Loveand Hadley (1986). Late Neogene and Quaternary geologyof the area between here and Stop 3 was first studied indetail by Wright (1946). For the next 25 mi the route continues to closely follow the Rio San Jose valley (via 1-40
and NM-6) through the easternmost Acoma—Zuni sectionof the Colorado Plateau to the western border of the Basinand Range province (Mexican Highlands section). See Love
and Young (1983) for a review of the geomorphic history ofthis part of the Rio Grande—Rio Puerco drainage basin.
Volcanic and extensional tectonic features along the route to
Stop 2 (Los Lunas volcano) are reviewed by Baldridge et al.
(1987, 1989).
From Stop 1 continue southeast on NM-279. 0.3 mi,roadcut in crest of hill through eolian facies of Morrison
Formation. Descend hill into valley past turnoff to Laguna
on right. 0.7 mi, cross Rio San Jose. Note younger valley fill
inset against Laguna basalt and dissection of alluvium bypresent stream. 1.0 mi, get into left lane to return to 140
East at Interchange 114. 1.9 mi, entering 1-40, dikes cutting
Morrison Formation exposed in roadcuts on both sides ofroad; descend hill through roadcuts of eolian facies ofMorrison Formation and Horse Mesa and Beclabito Members of Wanakah Formation cut by diabase dikes. At 4:00
dunes cascade down from mesa rim. 2.7 mi, wavy beddingin Wanakah Formation to right is due to collapse below. 3.1

122mi, clastic dikes in Horse Mesa Member of Wanakah on
both sides of 1-40. 3.4 mi, large collapse feature at 3:004:00 in Horse Mesa Member of Wanakah Formation; toeast across valley is section of Entrada Sandstone, Todilto
limestone and gypsum, Beclabito, and Horse Mesa
Members of Wanakah, and lower eolian facies of Morrison
Formation. Laguna flow is dissected in valley.
4.2 mi, MP 117, Mesita exit ahead; route is in Quaternary
alluvium and Triassic Chinle Formation. 7.2 mi, MP 120,bridge across Rio San Jose ahead. Valley at 4:00 is Arroyo
Colorado, location of several Pliocene and Pleistocene volcanic centers (Baldridge et al., 1987). 9.0 mi, railroad overpass; lowest cuesta on left for next several miles is Correo
Sandstone bed in Petrified Forest Member of Chinle Formation. Note toreva blocks partially buried by alluvium nearbase of slope to left. 10.2 mi, MP 123, section above Chinle
Formation exposed on Mesa Gigante to north consists ofEntrada, Wanakah, Morrison, and Dakota Formations, withthe latter forming top of mesa. 13.7mi, take exit 1-40 to NM6. Turn south toward Los Lunas (mi 14.0).
14.7 mi, MP 1 on NM-6, descend alluvial apron to Rio
San Jose. Mesa Redonda is at 10:00, Sierra Lucero at
12:00; Cerro Verde, volcanic source of the 0.32 Ma (K-Ar)
Suwanee basalt flow is at 1:00 (Bachman and Mehnert,
1978; Baldridge et al., 1987; Leavy and Shafiqullah, 1987).
The Suwanee basalt flow is 44 mi (70 km) long, flowing
north to the Rio San Jose, east, and then south. The flow ispartially covered with alluvium and eolian deposits andcontrols the amount of incision upstream from the flow.
15.0 mi, cross Rio San Jose. According to U.S. Geological
Survey records, present average discharge is 11.6 cfs (0.33
m3/ s). The largest estimated historic flood here was 11,000
cfs (312 m3/s) on August 21, 1935. 15.8 mi, route crosses
old US-66. Wild Horse Mesa Bar to left, Correo townsite
to west across railroad overpass.
17.9 mi, head of Rio San Jose Canyon in Suwanee
basalt to left; Suwanee Peak at 9:30; Mesa Redonda at
1:00-2:00 is composed of the same units exposed in Mesa
Gigante, from the Todilto gypsum up to the Dakota
Formation. The units have been folded and faulted and arecapped with a basalt flow about 3 Ma old. Beyond Mesa
Redonda is the northern edge of the Lucero uplift. The
northern end is capped with a basalt flow 3.4 Ma (K-Ar;
Bachman and Mehnert, 1978; Baldridge et al., 1987).
Numerous landslides have developed in the mudstones of
the Chinle Formation around the margin of the mesa. 18.7
mi, MP 5, site of Suwanee Station on AT&SF Railroad toright. 20.7 mi, roadcuts in Todilto Gypsum. To east aresqueeze-ups of Suwanee basalt. Mesa beyond San Jose
Canyon at 11:00 consists of Jurassic and Cretaceous
sandstones. There is no sharp topographic break between
the Colorado Plateau and Basin and Range physiographic
provinces at this point (Hawley, 1986a), but we are close
to the boundary shown on most physiographic maps (e.g.
Fenneman, 1931; Hunt, 1974).
22.5 mi, view at 9:30 down Rio San Jose toward Rio
Puerco. This area provides spectacular views of the northern
Lucero uplift and lower Rio San Jose. The Rio San Jose
joins the Rio Puerco about 3 mi (5 km) downstream. Notethe reddish color of the San Jose alluvium, derived from
Permian and Triassic redbeds. 23.2 mi, the structural boundary between the Colorado Plateau and the Rio Grande rift isbeneath the basalt flow in this area. The rift-bounding Santa
Fe fault along the front of the Lucero uplift has been

projected northward to 1-40 and farther north toward theNacimiento Mountains. Note Manzano (12:00) and Sandia(10:00) uplifts on the far eastern skyline beyond the Llanode Albuquerque. See Lozinsky et al. (this volume) for geological overview of the Albuquerque (rift) basin and Santa
Fe Group basin-fill stratigraphy.
25.7 mi, MP 12, Ladron Mountains are at 12:00, faulted
edge of Lucero uplift at 12:30-1:00; fill of Albuquerquebasin is exposed in hills east of the mountain front. Travertine from springs along the faulted edge of the Lucero uplift
makes up the white deposits at 3:00. Hidden Mountain at
10:30. 27.3 mi, road descends off tip of Suwanee (Cerro
Verde) basalt flow toward Arroyo Garcia. This is a site ofexperimental work on surface-exposure dating of basalt
flows using the CL-36 method (Phillips et al., 1986; Leavyet al., 1987). Note ancestral Rio San Jose gravel beneath theflow. An outlier of Santa Fe Group basin fill (late Miocene
to early Pleistocene) at 10:00 lies between the valley of theancestral Rio San Jose and Rio Puerco. 31.7 mi, MP 18,
valley of Rio Puerco at 10:00-1:00; Hidden Mountain at
2:30 is a subvolcanic igneous intrusion cutting basin-fillsediments. The plug is dated at about 8 Ma.
32.1 mi, bridge over Rio Puerco. Gaging station at
AT&SF Railroad bridge to right. According to Hawley,
Love, and Wells (1983), the average discharge is 57 ft3/sec(1.6 m3/ sec). Water flows here about 55% of the year (200days). Bankfull discharge for the present inner channel isapproximately 3250 ft3/sec (92 m3/sec). Extreme floods
occurred on August 12 and September 23 of 1929. The
flood of September 23 had a flood stage of 18 ft (5.5 m),
destroyed the railroad bridge, and had an estimateddischarge of 37,700 ft3/sec (1,070 m3/sec). The highest
recorded sediment concentration for the Rio Puerco was
680,000 ppm, 75% of which was sand (Nordin, 1963).
The Rio Puerco in this reach has developed an active
inner channel and inner floodplain within Rio Puerco arroyo
since the floods of 1929. Common geomorphic and sedimentologic features within the channel include thalweg,
sand bars, several levels of battures, and berms with ripples
and dunes. Features developed along the inner floodplain
include natural levees, chute channels across point bars,
oxbows, splays, and accumulations of eolian sand. Tamarix(Tamarix sp.), willows (Salix sp.), and rabbit brush (Crysothamnus nauseosus) dominate the vegetation within thearroyo. Also see discussion in Hawley, Love, and Wells(1982) and Wells et al. (1982, 1983). Stop 2-3 is about 13
mi (21 km) down valley from this point.
32.3 mi, ascending alluvial apron derived from basin fill.
The basin fill (upper Santa Fe Group) in this area was deposited by southeastward-flowing ancestral Rio Puerco-Rio
San Jose systems. Fossils within the fill near here are between 1.4 and 4 Ma (Blancan age). 32.7 mi, MP 19, Mohinas Mountain south of Hidden Mountain at 2:30 is anotherigneous intrusion cutting basin fill. For the next few miles,
tributary washes crossed by the road were quite active on
July 1, 1986, resulting in road damage and temporary closure. 34.3, mi, lower Rio Puerco valley at 2:00-3:30. Pottery
Mound Pueblo site and University of New Mexico researcharea on valley floor adjacent to Rio Puerco arroyo at 3:30.
The ruined 500-room pueblo was occupied between A.D.
1325 and 1475 and may have housed 1000 people (Cordell,1982a). Occupants brought pottery from all over theSouthwest to the site. Potsherds were thrown into a channelof the Rio Puerco about 20 ft (6 m) deep. The

channel filled during the occupation and the modern channels are north of the pueblo.
37.2 mi, ascending eroded edge of Llano de
Albuquerque; then crossing old valley graded toward Rio
Puerco. Cat Hills basalt flow from cinder cones at 10:30 has
been dated at approximately 140,000 yrs. 41.5 mi, bridgeover AT&SF Railroad. 41.6 mi, prepare to turn right. 42.4
mi, turn right on gravel road toward Cerro de Los Lunas(Los Lunas volcano). 42.7 mi, turn left onto dirt track thatheads southeast toward volcano along edge of Llano. Routeascends alluvial apron from Los Lunas volcano. 44.1 mi,turn left before juniper tree. 44.3 mi, park vehicles in wide
spot for Stop 2-2.
Stop 2-2: Los Lunas volcano (David W. Love, Richard P.
Lozinsky, and Bruce Hallett). This stop (elevation about 5500
ft, 1675 m) is on the northwest flank of Los Lunas volcano
along a 260 ft (80 m) escarpment exposing 660 ft (200 m) of
upwarped basin fill and a section of post-volcano alluvial
apron. The volcano and exposed basin fill lie along asouthwest—northeast-trending accommodation zone that
divides the Albuquerque basin into two half-grabens (Lozinsky et al., this volume). The northern Albuquerque basin is
deeper on the east side and steps up to the west; the southern
Albuquerque basin is deeper on the west side and steps up to
the east (see Lozinsky et al., this volume). Fifteen kilometers
to the north-northeast from this stop, a Shell Oil wellpenetrated 4407 m (14,458 ft) of upper Oligocene to lower
Pleistocene basin fill (Santa Fe Group) and did not reach preTertiary rocks at a total depth of 6482 m (21,266 ft). The
younger basin fill (Plio-Pleistocene Sierra Ladrones
Formation; Machette, 1988a, b) obscures structural differences between the two sub-basins. The Llano de Albuquerque,the nearly planar surface west of Los Lunas volcano

123that slopes gently to the south, is a remnant of a broad fluvial
plain that predated initial river-valley entrenchment in early
to middle(?) Pleistocene time. To the northwest, the Cat
Hills are a line of cinder cones capping the Llano de
Albuquerque. A related basalt flow that descends to a terrace(probably Segundo Alto Surface of Bryan, 1938) has a K—
Ar date of 140,000 ± 38,000 yrs (Kudo et al., 1977). To thenorth is Wind Mesa, an undated shield volcano of basaltic
andesite with a post-eruptive graben trending north—south
through the center. Farther north-northeast are the
Albuquerque Volcanoes, Sierra Nacimiento, and Jemez
Mountains. To the northeast is Isleta volcano, a tuff-ring,cinder cone, and flow complex of alkali-olivine basalt dated
at 2.78 ± 0.12 Ma (Kelley and Kudo, 1978), which is interbedded with ancestral Rio Grande deposits.
Los Lunas volcano consists of five flows, two vents, andsome intrusive centers (Kelley and Kudo, 1978). A
separate older, faulted sequence of flows occurs to thesouth and west (see below). Conventional K—Ar dates on
several of the flows on the main volcano range from 1.31 ±0.05 to 1.01 ±0.10 Ma (Bachman et al., 1975). The oldest
flows of Los Lunas volcano lie on an angular unconformity
on previously folded and possibly faulted Santa Fe beds
(Fig. 2-1).
The series of flows seen on the north side of Los Lunas
volcano are light-brown to gray in color, medium- to finegrained, and finely vesiculated. Quartz-rich blebs and granitic and mafic granulite xenoliths occur in some flows. A
1-6 m (3-20 ft) reddish-orange spatter-fed flow is seen

124
bedded between two darker flows. With the exception ofthis flow, others seen here in outcrop can be traced to theeast, descending into the Rio Grande valley. There is noevidence that these flows have been buried by Rio Grande
gravels. The intrusive andesites, observed mainly to thesouthwest of the central peak, are typically a darker gray to
black, finer-grained, and also finely vesiculated.
The volcanic rocks at Los Lunas define two chemicallydistinct trachyandesite groups. The intrusive rocks areslightly subalkalic, while the flows tend to be more alkalic incomposition. These intrusives, although more "mafic"
looking, have higher SiO 2 and lower TiO2, MgO, CaO,
Na 2 O, Sr, Zr, and Nb (Hallett, unpubl. data) compared to theflows. High K/P and Ba/Nb, low Y/Nb and Zr/Nb, positively
correlated "SeSr and SiO 2 values, in conjunction with crustal
xenoliths all suggest that the Los Lunas trachyandesites haveundergone crustal contamination.
Escarpments northwest and southwest of Los Lunas volcano expose similar basin-fill stratigraphy. The base of thesection directly below Stop 2-2 is a local angular unconformity between a deformed pumice-bearing fluvial unit
dipping from 12 to 85° to the northwest, and a nearly horizontal overlying alluvial channel and eolian complex. To
the north, this complex dips 8-15° northwest and accounts
for the bulk of the exposures seen at this stop. Lozinsky(1988) measured about 145 m (476 ft) of basin fill between
the angular unconformity and the base of tephra from Los
Lunas volcano. The lower part of the section is dominatedby pedogenically modified eolian sand with rare crossbedding. The upper part of the section is fluvial sandy gravels
from the northwest (Rio Puerco drainage). Within the section are several pumice- and obsidian-bearing gravels derived from the Jemez Mountains. One distinctive unit is astack of pumice debris flows within a thicker section of
overbank silts and clays (a date is pending). Blancan fossils
are reported from this section (Galusha, unpubl. data; Tedford, 1981), but none were found during the current study.
The lowest fallout tephra from Los Lunas volcano isabout 1 m thick where it is preserved in a syncline closest to
the volcano (1 km north of Stop 2-2) and is deformed (seeHaneberg, this volume). This tephra is reworked to fill the
syncline with horizontal debris flows. Within a smaller graben farther north, three or more fallout tephra beds drape
across eolian sand and clay dunes (Fig. 2-2). Farther northalong NM-6 the single tephra layer thins to 20 cm.

Overlying the tephra with a local unconformity at thebase is the undeformed alluvial apron shed from Los
Lunas volcano and uplifted basin fill. The alluvial apron
ranges up to about 50 m (165 ft) thick. At Stop 2-2, agravelly channel truncates the older basin fill. Ash fromthe 1.1 Ma old upper Bandelier eruption (Tsankawi tephra
at base of Tshirege Member, communication of Manley,
1976; referenced by Bachman and Mehnert, 1978) is wellpreserved in this channel fill along with boulders of Los
Lunas andesite. The upper part of the alluvial apron
contains a stack of two or more soils with stage II to III
pedogenic carbonate horizons.
The southwest escarpment (over the hill from Stop 2-2)
exposes similar units, but the lowest part of the section
consists of andesitic lava domes and flows buried by eolian
units. The emplacement of these eruptives may account foi
the earlier deformation seen near the base of both exposed
sections.
Controversies regarding the stratigraphy and geomorphology around Los Lunas volcano may be settled by analysis of the stratigraphy and geochemistry of the flows. As
summarized by Hawley (1978; Hawley, Love, and Wells,
1982), Bachman (Bachman et al., 1975; Bachman and Mehnert, 1978) interpreted the upper paleosol-soil sequence as
the culmination of aggradation of the Llano de Albuquerquesurface. He also correlated the dated Los Lunas andesitewith the outlying, largely buried andesite flows low in thesection. Thus Bachman interpreted the Los Lunas volcano
as a partially exhumed volcanic center and the Llano de
Albuquerque surface as being post-Los Lunas volcano. He
also demonstrated that the Llano de Albuquerque surface is
not the same surface as the Ortiz erosion surface describedby Bryan (1938).
In contrast, Kelley (Kelley et al., 1976; Kelley, 1977;
Kelley and Kudo, 1978) interpreted the eruption of Los
Lunas volcano as post-dating formation of the Llano de
Albuquerque surface with flows descending to a lower surface within the Rio Grande valley. Kelley correlated thesurface along truncated basin fill at Stop 2-2 with the Ortiz
erosion surface of Bryan (1938) and Bryan and McCann(1938). Kelley also interpreted the andesites within the section south of Los Lunas volcano as being intrusives.
The controversy may be resolved by taking interpretations
from both groups. The southwestern andesites are demonstrably flows interbedded with Santa Fe Group eolian
sediments, are chemically distinct (Kasten, 1977), and areolder than Los Lunas volcano. The volcano erupted from theuplifted edge of the Llano de Albuquerque and lavas flowed
to a lower level to the east and were not buried by basin fill.
The amount of tectonism involved in the differential
elevation of the volcano and the flows remains to be
determined. The Llano de Albuquerque surface west of Los
Lunas stabilized by 1.1 Ma and was buried by the alluvial
apron shed from Los Lunas volcano and underlying deformed sediments. Soils associated with the alluvial apron are
products of pedogenesis that started 1.1-0.5 Ma ago. The
Llano de Albuquerque and the alluvial apron are both muchyounger than the mid-Pliocene Ortiz surface.
Directions and comments en-route to Stop 2-3
Stop 2-3 is in the Rio Puerco valley west of Belen. From
Stop 2-2 return 1.9 mi to NM-6 and turn right (east) 4.0 mi
to 1-25. Route descends to Segundo Alto geomorphic
surface, here a terrace about 40 m (130 ft) above the Rio

125
Grande floodplain, formed on mid-Pleistocene fluvial deposits of the Los Duranes alluvium (Lambert, 1968). Turnright (south) on 1-25. To west are flows of Los Lunas volcano. Continue south 11.2 mi to Exit 191 (Sosimo Padilla
Blvd.). Turn right (west) on Sosimo Padilla Blvd.; ascendhill past exposures of Sierra Ladrones Formation (upper
Santa Fe Group) to Llano de Albuquerque. Study of thespatial variability of facies and permeability of these deposits is summarized by Lohmann et al. (this volume).
Route crosses the Llano rim 0.9 mi west of 1-25
Interchange 191. This is the site of Optional Stop 2-4.
Continue past New Mexico Travertine processing plant to
end of pavement at Belen International Airport (2 mi).
Continue 0.6 mi and turn right (west) on county road(slightly improved). Road west crosses several subdued
fault scarps before reaching the escarpment "Ceja del Rio
Puerco" 2.6 mi from turn. Descend Ceja toward Rio Puerco
and turn right 0.9 mi below the escarpment rim toward
abandoned windmill. Drive 0.4 mi off county road past
windmill and park. Walk to arroyo edge at meander bend.
Stop 2-3: Rio Puerco (David W. Love and John W.
Hawley). At this stop (elevation 4900 ft, 1495 m) we examine depositional environments of the modern Rio Puerco
and an exposure of facies in a buried arroyo of late Holocene
age, to determine the way it filled. Topics covered besides
specific features within the arroyo and their paleohydrologic
implications are (1) the thickness and age of basin and valley
fills of the southern Albuquerque basin, (2) spatial variabilityof permeability within depositional units (Stop 2-4), (3)terraces along the Rio Puerco, and (4) regional and perchedground water in relation to recharge and evapotranspiration
along the modern Rio Puerco. This part of the lower Rio
Puerco valley is the site of several recent geoarchaeologicaland environmental geochemical studies relating to proposed
dam construction and downstream effects of uranium mine—
mill operations (Love et al., 1982; Love and Young, 1983;
Popp et al., 1984, 1988).
Exposures along the lower Rio Puerco (Shepherd, 1976,
1987; Young, 1982; Love et al. 1982; Love and Young,
1983; Robinson and Love, 1986) show at least two
episodes of arroyo cuts-and-fills similar in size and shape
to the modern arroyo and other buried channels withdifferent facies (Fig. 2-3). At Stop 2-3 we can compare
similar facies of the modern Rio Puerco and a buried
channel to speculate about similar processes and conditionsin the past and what to expect in the future.
The lower Rio Puerco is a losing stream, that is, it loses
water to its bed. A flow of as much as 0.14 m3/s (5 cfs) at
the stream gage along NM-6 west of Los Lunas is lost to
seepage and evaporation before it reaches the stream gageat Bernardo (Heath, 1983; Heath and Love, 1986). In thisreach west of Belen, the permanent water table is 10-13 m(30-40 ft) below the base of the modern channel, but there
are seasonally fluctuating perched water tables on clay
layers within the alluvium above the permanent water table(Figs. 2-4, 2-5). These perched water lenses tend to berecharged from the Rio Puerco when it is flowing on loose
sand, and to be depleted by seepage to lower levels and bytranspiration by tamarisks and willows.
The channel shape of the lower Rio Puerco reflects repeated attempts at adjustment to extremely variable flows
and an extremely high sediment load. Flows much larger
than bankfull tend to scour the base of the channel andeven scour the floodplain locally. Smaller flows tend tobuild sand bars attached to both sides of the channel and to
aggrade the floor of the channel.
Within the buried arroyo exposed in the modern arroyo
wall, extensive floodplain deposits are scoured by both large
and small channels (Fig. 2-6). The large channel shows
multiple episodes of scour and fill before the whole arroyo
was filled in. The "last gasp" channel of the fill is redder,more poorly sorted, and is interpreted to be due to local
runoff from the surrounding valley sides (local sources forred clay and well rounded gravel). Apparently less and lessrunoff and sediment from upstream reached the lower Rio
Puerco as its arroyo filled up. Does this imply (1) a drying
climate, (2) increased retention of water upstream due tobetter vegetation cover, or (3) some other complex response(Schumm, 1977) mechanism upstream in the basin resulting
in decreased sediment and/or water delivery? Can we draw
any conclusions regarding how the present arroyo will fillin? Can we draw any conclusions about climatic cyclicity
with conditions even more arid than now?
Valley fill and terraces along the Rio Puerco reflect
longer-term cyclic adjustments in sediment storage anddelivery (Love, 1986). The valley fill in this reach of the
Rio Puerco is at least 135 ft (41 m) thick. Drill holes havenot encountered extensive coarse facies at the base orwithin the fill. The most obvious terrace along the Rio
Puerco is the Montano Flats surface 100-140 ft (30-42 m)
above the Rio Puerco west of Stop 3-3; other possible
terrace remnants are at 10 ft (3 m), 20-25 ft (6-8 m), 80100 ft (24-30 m), and 180-200 ft (55-61 m) (Fig. 2-7;
Love and Young, 1983).
Directions and comments en-route to
Optional Stop 2-4
From Stop 2-3 return to the county road and back (4.1 mi)across the Llano de Albuquerque to the rim of theescarpment forming the western wall of the AlbuquerqueBelen Valley (elevation about 5175 ft, 1575 m). Turn righton dirt road and continue 0.5 mi east-southeast to tip of asmall mesa promontory (Stop 2-4). Note strong calcic soil
capping the Llano surface. The eastern rim of the Llano de
Albuquerque was designated Cejita Blanca (little white eye-brow) by Bryan and McCann (1938) because of the prominent
light-colored soil-carbonate horizon at the upper edge of theescarpment.
Stop 2-4 (optional): Hydrogeologic investigations of Sierra
Ladrones Formation. See Lohmann et al. (this volume) for
tour-stop discussion.
Directions and comments en-route to Stop 2-5
From Stop 2-4 retrace route on county road down theescarpment to the 1-25 Interchange (0.9 mi). Entering Belen;continue east on Sosimo Padilla Blvd. under the 1-25 overpass
to Main St. (1.0 mi). Turn left (north) on Main St. andgo 0.6 mi before turning right on Reinken St. (NM-309).
Continue 2.3 mi east on NM-309. After crossing a railroad
overpass (0.5 mi) and the Rio Grande (2.2 mi), turn righton NM-47 at Rio Communities. The route continues southeast 19.4 mi on NM-47 to its junction with US-60 (Stop 25). A New Mexico Geological Society Field Conferenceroad log by Hawley, Love, and Lambert (1982) covers thetrip route from Belen to US-60 (reverse orientation).
The Rio Grande floodplain at Belen is underlain by afining-upward valley-fill sequence of late Quaternary age
that is as much as 120 ft (36 m) thick (Titus, 1963). Most
large-capacity irrigation wells obtain water from a basal
fluvial gravel zone below 80 ft (24 m). Underlying sandy

128
fluvial deposits of an ancestral Rio Grande system (upper
Santa Fe Group, Pliocene) are the main producers for municipal and domestic water supplies.
For the first 6 mi on NM-47 (Rio Communities to turnoff
to John F. Kennedy Campground) the route ascends a sequence of fluvial terraces formed during middle to late Pleistocene episodes of valley cutting and partial backfilling(Titus, 1963). The highest surfaces (170-280 ft, 55-85 m) areoffset by north-trending faults (Machette, 1982; Machetteand McGimsey, 1983). Thin terrace fills, with calcic soils,
are inset into intertonguing fluvial and distal-piedmont faciesof the upper member of the Sierra Ladrones Formation (upper Santa Fe Group). Southeast of MP 11 the route ascends
to the broad piedmont alluvial plain (bajada) formed bycoalescence of fans at the western base of the Manzano
Mountains (maximum elevation 10,098 ft, 3079 m). The
major piedmont geomorphic surface dates from late middle
Pleistocene and has been designated the Llano de Manzano
by Machette (1978c, 1985). This surface is cut by the broad
valley of Abo Arroyo (MP 2). Between MP 9 and Stop 25
the route crosses shallowly buried segments of the two major
rift-boundary fault zones in the southeastern Albuquerquebasin, the Hubbell Springs and Manzano—Los Pinos faults(Kelley, 1977; Machette and McGimsey, 1983). Park at US60—NM-47 junction opposite MP 183.
Stop 2-5: Overview of southeastern Albuquerque basin
and Sevilleta National Wildlife Refuge area. The low
bedrock knob of Precambrian metasediments and metavolcanics (elevation 5700 ft, 1800 m) 250 m south of the US60—NM-47 junction offers a good view of the southeastern
Albuquerque basin and the Abo Pass area between the southern Manzano and northern Los Pinos Mountains (Fig. 2-8).
The complex Precambrian metamorphic terrane of the Los
Pinos block (Myers et al., 1981; Reiche, 1949; Stark andDapples, 1946; Stark, 1956) is thrust against upper Paleozoic
marine rocks along a high-to-low-angle reverse fault zone,
the early Tertiary "Montosa thrust" (Fig. 2-9). The lateCenozoic Los Pinos fault zone is buried by upper Qua-

ternary fan alluvium about 0.35 mi (560 m) west-northwestof this stop (Machette and McGimsey, 1983). The Consortium for Continental Reflection Profiling (COCORP) has
done deep seismic-reflection profiling across the Rio Grande
rift in this area (Brown et al., 1980; de Voogd et al., 1986).
Stop 2-5 is near the northeast corner of the Sevilleta
National Wildlife Refuge administered by the U.S. Fish andWildlife Service. The refuge (area about 400 me, 1036 km2)
is the center for a number of research projects involving bothbiological and earth sciences, including a National Science
Foundation-funded Long-Term Ecological Research Project(LTER). Paleoecological aspects of ongoing project research
are discussed below.

Vegetation history of the Sevilleta LTER
Julio L. Betancourt', Raymond M. Turner',and Steve Tharnstrom2'U.S. Geological Survey, Tucson, Arizona 84701-1393;'Sandia National Laboratories, Albu q uer q ue, New Mexico 87115

From south to north in New Mexico, creosote bush (Larrea tridentata) gradually gives way to one-seed juniper (Juniperus monosperma) as the dominant shrub in the lowlands.
This transition is most evident in and near the Sevilleta LongTerm Ecological Research (LTER) site, along the western
piedmont of the Los Pinos Mountains, where minimum
elevations rise above 5000 ft (1525 m). Here aggregates ofcreosote bush surround numerous snags of junipers that diedduring the prolonged drought of 1946-1950, a graphic
example of the dynamic interactions across this ecotone.
Knowledge about the past distribution of creosote bush,about the rates of its migration and subsequent expansion,
may be critical for anticipating its penetration to Albuquerqueand points beyond. Toward this goal, we have initiated astudy of the vegetation history of the Sevilleta LTER, based
primarily on packrat midden analysis and repeat
photography. Plant debris accumulated by packrats fromclose range (mostly <30 m away) commonly becomes encased in crystallized urine to form middens. These deposits
can persevere for tens of millennia in caves and rock crevices
and are now the primary source of vegetation reconstructionsin the arid interior of North America (Betancourt et al.,1990). Repeat photography involves comparison of vegetation in photographs taken at various times from the same
camera station (Hastings and Turner, 1965). The initial phaseof the study involves analyses of 50 middens from the Los
Pinos Mountains and numerous paired photographs from thearea in and around the Sevilleta LTER.
During the latest Wisconsin, woodlands of Colorado piñon(Pinus edulis) and Rocky Mountain juniper (Juniperus scopulorum) dominated limestone cliffs below 1830 m elevation.
A midden in Palo Duro Canyon (23 km south-southwest of
Stop 2-4) represents the northernmost glacial occurrence ofpiñon in New Mexico and Rocky Mountain juniper no longer
grows in the Los Pinos Mountains. Piñon needles in this
12,000 years old deposit contain high-molecularweight DNA;application of new molecular techniques, specificallypolymerase chain reaction (PCR), will help document changes
in genotype frequencies associated with piñon's migration
upward and northward during the Holocene. Piñon is now
missing from sites it occupied during the late Holocene
between 1770 and 1860 m. Piñon's lower limit retreated
upward about 100 m sometime during the last 300 years.
Packrat middens from Mesa del Yeso, at the southern

end of the Sevilleta LTER, document more widespread ocotillo (Fouquieria splendens) during the late Holocene, including north-facing limestone slopes. Today stands ofocotillo are restricted to south-facing slopes at its
northernmost limits in the Sevilleta LTER.
Accelerated expansion of creosote bush in central New
Mexico, poorly dated and sparsely documented, has been
blamed on overgrazing and droughts of the past two centuries (Gardner, 1951). In the long term, this expansion
represents the northern terminus of a migration that beganat the end of the last ice age. For reference, not until 4000
years ago did creosote bush colonize sites in far west Texasand southern New Mexico (Van Devender, 1990). The arrival of creosote bush on limestone slopes in the Sevilleta
LTER by 2500 years B.P. predates overstocking of ranges,
but colonization of fine-grained soils and other substrates(sandstone, granite, and basalt) happened only within thelast 200 years.
Directions and comments en-route to Mountainair
From Stop 2-5 (MP 183) continue 48 mi east on US-60
through Abo (15 mi), Mountainair (22 mi), and Willard (33mi) to Lake Estancia Stop E-1 near Lucy Siding (ATSF
Railroad). See Hawley, Foster et al. (1982) for moredetailed information on the tour route to Mountainair(reverse direction). Geologic maps by Myers (1977) and
Myers et al. (1981) also cover the first 11 mi (to MP 194).
Near MP 185, US-60 crosses the Montosa fault zone andthe edge of the major structural uplift (Manzano—Los Pinos)that forms the eastern boundary of the Rio Grande rift (Fig.
2-9). To the east is the transition zone between the Basin andRange and Great Plains structural provinces, which ischaracterized by gently east-tilted sedimentary rocks of latePaleozoic age. Pennsylvanian and Permian marine carbonates and Permian continental to marine clastic and gypsiferous evaporites dominate the section (Madera Group and
Bursum, Abo, Yeso, Glorieta, and San Andres Formations,in ascending order). The route follows the broad upper valley of Abo Arroyo. Extensive upper Quaternary valley-filldeposits have been deeply incised by modern arroyo channels. This is another area of New Mexico where work needs
to be done on valley cut-and-fill history. The Abo Ruins
Unit of Salinas National Monument (pre-1680 pueblo andmission) is located about 0.5 mi north of US-60 at MP 195.
MP 204. At the western edge of Mountainair, the routecrosses the divide (elevation about 6550 ft, 2000 m)
between the Rio Grande (Abo Arroyo) watershed and theclosed Estancia drainage basin (Fig. 2). This point is also
near the boundary between the Mexican Highland (RioGrande subsection) and Sacramento sections of the Basinand Range physiographic province (Fenneman, 1931;
Hunt, 1974; Hawley, 1986a). The Sacramento section
comprises a physiographic transition zone, here about 50
mi (80 km) wide and characterized by high tablelands andshallow structural basins, that lies between the Pecos valley
section of the Great Plains and the high fault-block ranges
forming the eastern edge of the southern Rio Grande rift.
Solution-subsidence landforms are locally a major featureof the Sacramento section.

130
Day 2, Part 2: Lake Estancia basin tour
(Leaders: Bruce D. Allen and John W. Hawley)

Introduction
For about 10 mi east of Mountainair (from MP 205) theroute descends across a broad, moderately dissected piedmont slope which grades to a base level slightly above thehighest shorelines of Pleistocene Lake Estancia (about 6210
ft, 1893 m; Fig. 2). Soils on the broad interfluves (Bourlier
et al., 1970) have well-developed calcic horizons that are
locally strongly indurated (morphogenetic stages III to V ofGile et al., 1966; and Machette, 1985). Basin-fill thickness
increases eastward from about 100 ft (30 m) near Mountainair to about 250 ft (75 m) along the basin axis near
Willard (Fig. E-1; Smith, 1957). The basin fill includes
alluvial, eolian, and lacustrine facies of Plio-Pleistocene
age, and the oldest deposits may correlate with the upper
Miocene Ogallala Formation of the southern Great Plains
region (Titus, 1973; Frye et al., 1982; Hawley, 1984).
Lower Permian Abo Formation redbeds (mainly sandstone
and mudstone) are the dominant substrate west of the highEstancia shoreline, while interbedded sandstone, siltstone,
gypsite, and limestone of the Yeso Formation immediately
underlie basin fill in much of the area inundated by Lake
Estancia (Smith, 1957; Titus, 1969, 1973). Upper Permian
limestone, gypsite, and sandstone of the San Andres andGlorieta Formations cap the high plateau (Chupadera andJumanes Mesas) that forms the south rim of Estancia Valley(Bates et al., 1947; Hawley, 1986b). The following comments cover the tour route to Stops E-1 and E-2 in thesouthern Lake Estancia basin. Stop E-3 is discussed in Part
1 of Day 3.
Directions and comments en-route to Stops E-1
and E-2
MP 215 (10 mi east of Mountainair). Scarp ahead (elevation 6200-6220 ft, 1890-1896 m) is a complex erosionalfeature initially cut during highest stands of middle(?) to late
Pleistocene Lake Estancia (Meinzer, 1911). No clear-cut
evidence has yet been found for shoreline features higher
than about 6210 ft (1893 m). Allen (this volume) describes

ongoing University of New Mexico studies of Late Quaternary features of the ancient lake-plain area (lake, playa,and dune deposits and associated wave-cut and deflational
landforms).
One mile east of wave-cut scarp is the junction with NM41 (start of Day 3 route to Stop E-3). Continue east on US60 through Willard (MP 218; elevation 6100 ft, 1860 m) on
nearly level basin floor, here a relict Wisconsin lake bed thathas been modified by late Quaternary eolian processes.
About 0.7 mi east of the NM-42 junction in Willard (rounding southeast bend in highway) clay dunes (lunettes) from9:00 to 2:00 are associated with lake-plain deflation.
East of MP 222 (rounding northeast bend in highway) theroute crosses southern end of Laguna del Perro, a northtrending playa-lake depression that is about 12 mi (20 km)
long and about 1 mi (1.6 km) wide (Fig. 2). This deflation
basin is the largest of nearly 100 such features that occur on
the valley floor. Several deflation basins with playa lakes willbe crossed in the next 4 mi. Accumulation of deflatedmaterial along downward margin of the basins has resulted inlarge clay dunes (lunettes). Near MP 224 note small mounds
to the left (north) forming hummocky topography. These
gypsum-rich dunes (Titus, 1969, 1973) predate the major
deflational basins, such as Laguna del Perro, and date fromthe initial desiccation of Wisconsin Lake Estancia 12,000 to
11,000 yrs ago. MP 225. Large deflation basins with playa
lakes to left and right. Stop E-2 will be at rest area ahead onleft. MP 226. Crossing another area dotted with low,
gypsum-rich dune mounds. 0.8 mi east of MP 227 crossnorth-trending gypsum beach ridge (elevation about 6100 ft,
1860 m) described by Allen (this volume, fig. 1) anddiscussed at Stop E-1. MP 231, on highest late Wisconsin
shoreline of Lake Estancia opposite railroad cut in Permian
Glorieta Sandstone at Lucy Siding (ATSF Railroad; elevation
6190 ft, 1887 m). Turn around, retrace route to MP 230, andpark near windmill on north side of US-60.
Stop E-1: Lake Estancia shoreline. This stop (elevation

13]
about 6145 ft, 1873 m) at the southwest end of Lucy Siding
is on one of the most prominent shoreline features of Meinzer 's (1911) Lake Estancia (Fig. 2). Landforms include
beaches, wave-cut benches, spits, and bay-mouth bars between 6090 and 6210 ft (1856 to 1893 m). The shorelines
here are in Harbour's (1958) "Group C" and appear as
undulations in the highway to the northeast and southwestof this stop. Lacustrine features at and below an elevation of
about 6200 ft (1890 m) are associated with latest Pleistocene (late Wisconsin), glacial-maxima highstands. At the6200 ft (1890 m) level, Lake Estancia would haveinundated an area of about 430 mil (1114 km2) and was over
140 ft (43 m) deep. This highstand occurred between about
15,500 and 19,500 yrs B.P. (Allen, this volume). The linear
feature about 1 mi (1.6 km) to the west-northwest, justabove the basin floor, is the gypsum beach ridge noted eastof MP 227. The reworked gypsum in the beach ridge wasderived from the small gypsum dune mounds that dot therelict lake-plain surface to the west at elevations below
6100 ft (1860 m). The gypsum beach ridge corresponds to a
25 ft (8 m) deep lake that postdates initial (pre-Holocene)desiccation of Lake Estancia after a major highstand about
13,000 yrs B.P. (Allen, this volume, fig. 4).
The lowest point on the Estancia drainage-basin divide(elevation 6340-6345 ft, 1932-1933 m) is located near Cedarvale, about 19 mi (30 km) south-southeast of this stop.
The surficial deposits in the divide are conglomeratic sandstones of Pliocene to latest Miocene(?) age, correlated by
Frye et al. (1982) with the Ogallala Formation. Surficial
beds are cemented into a pedogenic calcrete (caprock caliche)
layer
with
dense
pisolitic
morphology(morphogenetic stage VI of Machette, 1985; McGrath andHawley, 1987). There is no evidence of Quaternary lake or
fluvial discharge over this "sill" into the chain of large
structural and/or solution-subsidence basins (Pinos Wells,
Encino, Vaughn) located to the northeast.
Evidence for and against Pleistocene surface drainage fromthe Estancia basin, across the east-bounding and basementcored Pedernal uplift, and onto the Pecos slope has beenpresented by a number of workers (notably Lyons, 1969;
Titus, 1969, 1973; Bachhuber, 1971, 1982, 1989; and Kelley,1972). The presence of abundant trout fossils in late
Pleistocene lake sediments (Bachhuber, 1989) indicates that apast surface drainage connection has existed, but not necessarily via the Encino—Pintado Canyon spillway route suggested by Bachhuber (1971, 1982) or the Encino—Vaughnroute discussed by Kelley (1972). Trout could have been
introduced through ancestral stream systems (present Rio
Grande or upper Pecos tributaries) that headed in the southern
Sangre de Cristo Range and originally discharged into the
Estancia basin in Pliocene or early Pleistocene time.
Perennial streams and springs in the higher parts of the
Manzano Mountains could have provided refugia for fish
populations during dry intervals after diversion of Sangre deCristo tributaries to external drainage systems. The importantrole that subsurface-flow systems may play in Lake Estancia
paleohydrology is discussed at Stop E-2.
Return on US-60 to rest area 0.4 mi southwest of MP 226
and Stop E-2; cross gypsum beach ridge 0.2 mi southwestof MP 228.
Stop E-2: Lake Estancia stratigraphic record. Details of
the lacustrine stratigraphic record exposed in deflation basins near this stop are presented by Allen (this volume).

Note the tan and gray lake beds exposed in the northern
escarpment of the playa-lake basin opposite the rest area.
These sediments of latest Pleistocene age (<20,000 yrs) areoverlain by a white horizon marking the top of the lacustrine
sequence and the position of the ancient valley floor (constructional lake-plain surface). Overlying sediments are
eolian deposits associated with deflation of the large playalake depressions of Holocene age.
The section exposed here is only the upper part of asequence of lake, playa, eolian, and minor alluvial beds
described in detail by Titus (1969, 1973) and Bachhuber(1971, 1989) at a number of surface and borehole sites in
the lake-plain and deflation-basin area north of this stop.
Sediments associated with middle to late Pleistocene episodes of lake expansion (Dog Lake Formation of Titus,
1973), including an early late Pleistocene interlacustral interval, have a maximum thickness of about 100 ft (30 m).
Underlying sand, clay, and pebble-gravel deposits with
gypsum- and carbonate-impregnated zones are part of an
extensive pre-late Pleistocene deposit (Estancia Formationof Titus, 1973) that rarely exceeds thicknesses of 300 ft (90m; Smith, 1957). Titus (1973) suggested that much of thisolder fill was deposited by an ancestral through-going drainage system which was disrupted by structural downdroppingof the present valley area relative to the bounding Manzano,
Sandia, and Pedernal uplifts. The possible significant role ofsolution subsidence resulting from partial dissolution of
gypsum beds in the underlying Yeso Formation and basin
fill has been recognized, but is not well documented.
Regional and local ground-water flow systems in the Estancia basin have been studied by Meinzer (1911), Smith(1957), Titus (1969, 1973, 1980), DeBrine (1971), andJenkins (1982). Most of these workers have assumed thatthe basin had been closed in terms of both surface andshallow subsurface flow during much of middle to late Quaternary time (Brady et al., 1984). However, some (startingwith Meinzer) have inferred that a deep interbasin-flow regime may exist, at least with respect to the major carbonate
aquifer in the region, the Pennsylvanian Madera Group(Jenkins, 1982, pp. 383-384).
Recent work by Hawley (1986b) on a hazardous-material
disposal site on Chupadera Mesa shows that contrary toother published information, there is no major ground-water
divide on Chupadera and Jumanes Mesas that would limit
subsurface discharge from Estancia basin during the documented lake maxima (as high as 6210 ft, 1893 m) since late
middle Pleistocene time (last 150,000 yrs). This study alsosuggests that a broad, south-trending fault zone (complex
extensional, compressional, and strike-slip components with
pre-Permian initial deformation), extending the length of
Chupadera Mesa, provides a connection between the regional carbonate aquifer and the shallow Estancia basin fill
and Yeso Formation ground-water system. Thus, there may
have been times of significant discharge from the Estancia
hydrologic basin southward into the Tularosa basin, the siteof late Pleistocene Lake Otero of Herrick (1904) and Meinzer and Hare (1915).
Recent studies on a candidate site for the Superconducting
Super Collider (SSC) project in northern Estancia basin(Johnpeer et al. , 1987) also support the hypothesis (first
proposed by Bates et al., 1947) that the north—south zone ofstructural deformation on Chupadera Mesa extends northward through the present Estancia basin into north-central
New Mexico (see Woodward et al., 1978). Laguna del Perro132
is aligned along this trend, as is a Laramide fault zone that
affects the Glorieta—Yeso aquifer system in the northeastern
Estancia Valley. In addition, detailed stratigraphic and geotechnical studies during the two-year SSC project revealed nolacustrine deposits or landforms at elevations above 6220 ft(1896 m) in the Lake Estancia basin.
A question that needs to be asked at this tour stop concerns
the entire problem of lake-basin paleohydrology, including therole that the playa-lake basins have played throughout middle
to late Quaternary time, both in terms of groundwater outflow
as well as inflow, and in the context of both regional and local
aquifer systems. Paleohydrologic models(Meinzer, 1922; Leopold, 1951; Antevs, 1954; Galloway,
1970; Brakenridge, 1978; Smith and Anderson, 1982) thatneglect ground-water—surface-water linkages cannot explain
many of the lacustrine and playa-basin features observed in
the Estancia basin. This is a very exciting area with great
potential for research in the Quaternary sciences (see Woodand Sanford, 1990; Kreider et al., 1990). The excellent fieldstudies by workers such as Meinzer, Smith, Harbour, Lyons,
Titus, and Bachhuber provide a solid basis for all current andfuture investigations.
Return to Mountainair and historic Shaffer Hotel via
US-60; end of Day 2.

Day 3, Part 1: Continuation of Lake Estancia basin tour
(Leaders: Bruce D. Allen and John W. Hawley)
Directions and comments en-route to Stop E-3
Return on US-60 to junction with NM-41. Continue 18.6
mi north on NM-41 through Estancia (10.6 mi) and McIntosh (18.2 mi) to junction with Torrance County Road
A72 (0.6 mi north of MP 18 and 0.4 mi north of McIntosh
Post Office). Turn left on CR-A72, and continue west 1.5
mi to crossroad. Turn left (south) and follow section-line
road 0.9 mi to junction with unimproved road to gravel pit
in lower Cienega Draw. Turn right (west) and continue
about 0.7 mi up the draw to floor of gravel pit and Stop E-3(Fig. 2).
Refer to Day 2 comments on the Lake Estancia basin
tour from Mountainair to the NM-41 junction west of
Willard. The route north on NM-41 through Estancia and
McIntosh is almost entirely on a relict-lake-floor surface(elevation range 6100-6140 ft, 1859-1871 m), which dates
from the last major highstand of Lake Estancia (about
13,000 yrs B.P.). Fig. E-2 is an oblique aerial photo ofshorelines between 6150 and 6180 ft (1875-1884 m), about
1 mi southwest of Estancia. About 0.8 mi west of NM-41(elevation 6170 ft, 1881 m) County Road A72 crosses
ridges and swales of the Wisconsin glacial-maximum
shoreline. The elevation range of nearshore gravel deposits
at Stop E-3 is about 6185 to 6200 ft (1885-1890 m).

Stop E-3: Lake Estancia glacial-maximum shoreline de•posits. Nearshore deposits associated with the glacial-maximum highstand of Lake Estancia are exposed in the wallsof numerous gravel pits along the western margin of theancient lake basin. Gravel-size clasts consist primarily of
Madera Limestone, which forms the eastern slope of the
Manzano Mountains to the west. At this locality, the upper
part of the shoreline sequence exhibits several decimeterscale cycles which consist of a basal, planar erosion surface
overlain by fine-grained sediments which coarsen upward
into landward-dipping, parallel-bedded gravels. The bedded
gravels at the top of each cycle are abruptly overlain by
fine-grained sediments of the next cycle. One interpretation
for this succession of asymmetric, coarsening-upward
cycles is that they represent transgressive—regressive
cycles of the ancient lake, with an erosional transgressive
surface occurring between the gravel beds and the overlying
fine-grained sediments of the next cycle.
The late Pleistocene nearshore deposits are overlain by aneolian "cap" of fine-grained material with an associatedweakly developed, cummulic soil that extends into thebeach gravels. The soil has developed since Lake Estancia
fell below this elevation, or subsequent to the last glacial
maximum. Subsurface soil horizons are light in color due
to illuvial carbonate, and exhibit stage II carbonate
morphology (after Gile et al., 1966).
Directions and comments en-route to 1-40 at Moriarty
Retrace route to NM-41. Turn left and continue 9.2 minorth on NM-41 to Moriarty and the junction with the business route of 1-40 (historic Highway 66).
For 5.5 mi north of the Stop E-1 turnoff (County Road
A72) NM-41 rises gradually (from 6150 to 6200 ft, 1875 to1890 m) across the gently undulating topography constructed during the Wisconsin glacial-maximum stage ofLake Estancia. A large gravel-bar complex deposited at themouth of Buffalo Draw is located about 500 yds east of
NM-41 at the 5.5 mi point. About 0.5 mi farther north, oneof the highest remnants of beach deposits preserved in thebasin is exposed in an old railroad cut 200 ft east of the
highway (elevation about 6210 ft, 1893 m). This outcrop of
partly cemented pebble gravel and coarse sand may be

133a shoreline deposit associated with a glacial maximum predating the last interglacial (>130,000 yrs). Exposures of suchpossible late middle Pleistocene lake deposits are very rare in
the Estancia Valley because deep dissection of the basin filloccurs mainly in the deflation basins of Laguna del Perroarea (Stops E-1 and E-2). Prime lake-front property ahead in
Moriarty during the next glacial maximum.
End of Lake Estancia basin tour.
Directions and comments en-route from Moriarty
to Albuquerque
From Moriarty continue west on 1-40 about 22 mi to the
Tijeras exit and then take old Hwy 66 (NM-333) throughTijeras Canyon (6.8 mi) to the eastern edge of Albuquerque atCentral Ave. and Tramway Blvd. The zero mileage point forthis trip segment is the NM-41 and Business 1-40 junction
near the north edge of Moriarty (Fig. 2).
Go west on Business 1-40 to overpass and westboundonramp of 1-40 (1.7 mi). Route ascends on lower and middle
Pleistocene piedmont deposits derived from Pennsylvanianand Permian rocks of the Manzano and Manzanita Mountains(Reiche, 1949; Titus, 1980). Engineering geology in this area
was investigated by Johnpeer et al. (1987) as part of theSuperconducting Super Collider project (SSC). Depth to
bedrock (Permian Abo and Yeso Formations) near Moriarty
is about 150-250 ft (46-76 m). Depth to water is currently
about 50-100 ft (15-30 m). MP 193. 3.2 mi, ascend from
Quaternary alluvium onto older basin fill (QTa) of late(?)
Miocene to early(?) Pleistocene age that may be partly
correlative with the Ogallala Formation. 5.9 mi, the New
Mexico version of the SSC was to pass beneath 1-40 in thisarea in a tunnel 130 ft (40 m) below the surface. 7.5 mi,outcrops of Upper Pennsylvanian Madera Group seen inarroyo and in hills to southwest. For next 3 mi, the area northof the highway is underlain by caves and karst.
8.5 mi, Edgewood exit; young alluvium in valley; older

basin fill exposed in low ridges to north; Madera Group to
south and west. Fig. 3-1 is a hydrogeologic section (Titus,
1980) covering the tour-route area between here and the
Sandia Mountain crest on the horizon. At 1:30-2:30 is South
Mountain, a mid-Tertiary monzonite laccolith with Permian
Glorieta Sandstone and San Andres Limestone on top. Beyond South Mountain at 2:30 are the San Pedro Mountains,another mid-Tertiary intrusive with abundant gold and copper mineralization. 9.1 mi, Madera Group outcrops before
crossing young alluvium in Juan Tomas Canyon; entering
area of Pennsylvanian—Permian rocks mapped by Myersand McKay (1976). 13.7 mi, Sedillo exit; drainage divide ofEstancia basin at crest of hill. Descend into Tijeras Canyon
cut in Wild Cow Formation (upper unit of Madera Group).
16.4 mi, fault in Wild Cow Formation down to northwest.
Northwest-dipping section on right for next 0.3 mi (presently covered with wire mesh and concrete) originally
showed cyclic deposition (clastic to carbonate) in membersof Wild Cow Formation below the Permian Abo Formation.
16.6 mi, road curves southwest into strike valley of Permian
Abo Formation. Permian Yeso Formation forms ridge tonorthwest of 1-40 for next mile.
17.4 mi, Zuzax exit. 18.3 mi, road passes northwestdipping Permian Yeso, San Andres, sliver of Triassic
Chinle, and Tertiary lamprophic dikes before crossing the
Gutierrez fault, a major branch of the Tijeras fault that
forms the southeast side of a graben holding the Tijeras
syncline and anticline seen topographically expressed inresistant Cretaceous Mesa Verde Sandstone to the north onridge. 18.7 mi, road crosses alluvium burying Gutierrez
fault and returns to Yeso and Abo Formations. Road
continues southwest in strike valley of Abo Formation on
local reddish-brown alluvium.
20.0 mi, take exit to NM-14. At stop sign turn left (south).
20.3 mi, underpass (below 1-40) and 90° bend to right inroad. 20.8 mi, stop sign. Village of Tijeras. Continue straight

134
through intersection on US-66. Village of Tijeras wasfounded in 1856 near site of a prehistoric Indian Pueblo(Cordell, 1982b). See Connolly et al. (1982) for moredetailed information on the tour route from Tijeras to
Albuquerque (reverse direction). The complex of canyons
crossing the mountains resembles open scissors, thetranslation of "tijeras." 21.0 mi, roadcuts are Lower Permian
Abo Formation. 21.8 mi, good exposures of highly deformed
Precambrian metavolcanics of Tijeras greenstone in roadcut
to left under boulder gravels.
22.1 mi, Seven Springs to left. Both alluvium and fractured
Precambrian rocks serve as aquifers (Titus, 1980). Alluviumin the inner valley locally is more than 100 ft thick. South of
Tijeras is faulted section of Sandia Formation and lower
Madera Group overlying Tijeras greenstone; north of
highway is a sliver of greenstone on south side of Tijeras
fault, a major fault showing both strike-slip and later verticalmotion. The highway crosses the fault at an oblique angle
between here and Deadman's curve. North of the Tijeras fault
is meta-arkosic Cibola gneiss, isoclinally folded quartzite, andSandia granite. Radiometric ages for the Sandia granite areabout 1445 ± 40 Ma and for the Cibola gneiss about 1576 ±73 Ma. Tertiary lamprophyre dikes cut the gneiss and granite
down canyon. Smith et al. (1982) discuss the stepped andstacked sequences of Quaternary alluvium and colluvium andassociated paleosols that cap canyon and piedmont erosion
surfaces from here to the mouth of Tijeras

Canyon. 22.4 mi, terrace gravel and fan deposits to left cap
metamorphic rocks in cut above Tijeras Creek.
22.9 mi, Deadman's curve cut through Cibola gneiss andquartzite. 23.4 mi, gradational contact between Cibola gneiss
and less and less foliated Sandia granite in slopes to right.
Note spheroidal weathering of granite. The "granite" in thisarea is generally quartz monzonite to granodiorite with large
pinkish porphyroblasts of microcline of metasomatic origin.
Differential weathering makes mafic inclusions stick out on
the granite corestones like raisins out of cookie dough. 24.1
mi, Village of Carnuel and Tijeras Creek to left. This old
Spanish village was established in the 18th century, but was
repeatedly destroyed by Apache raids. To the south, gold,
barite, and fluorite mines were developed in Precambriangreenstone terrane. Route ahead on pediment cut on Sandia
granite with discontinuous veneer of Quaternary alluviumand colluvium deposited in apex area of Tijeras Creek fan.
25.2 mi, onramp to 1-40 to right. Bear left and cross overpass. 26.4 mi, break in slope (piedmont angle at 3:00 to 4:00)
between mountain front and piedmont erosion surfaces on
Sandia granite at western end of Tijeras Canyon. 27.0 mi,
sharp curve to left. Crossing shallowly buried segment of
Sandia fault zone (Kelley and Northrop, 1975; Machette,1982). 27.1 mi, stop sign. Bear right. 27.3 mi, intersection of
Tramway Blvd. and Central Ave. and start of Albuquerque—
Rio Rancho urban-area tour.

Day 3, Part 2: Albuquerque—Rio Rancho urban-area tour
(Leaders: John W. Hawley and Stephen G. Wells)
SummaryThis tour of the Albuquerque area (Fig. 2) emphasizes
environmental and engineering geology, geomorphology, andQuaternary stratigraphy of valley- and basin-fill deposits. The
route starts at the Central Ave. (Hwy 66)—Tramway Blvd.
intersection and traverses the urban area in a loop about 58
mi long. Local points of interest are discussed at four stops.
Much of the following discussion has been adapted from roadlogs by Hawley, Lambert et al. , Kelley et al. , and Lambert etal. (1982). Fig. A-1 is a diagrammatic cross-section of thenorthern Albuquerque basin that emphasizes late Cenozoic
features (landforms, basin and valley fills, ground water, andmajor structural components) in the area between Stops A-1,
A-3, and A-4.
The first segment of the tour route, via Tramway Blvd.,
crosses the upper piedmont alluvial plain (Llano de Sandia)
at the foot of the Sandia Mountains. Stop A-1 is at Elena
Gallegos Municipal Park on upper Arroyo del Pino. This
stop provides an overview of the northeast Albuquerquearea and western front of the Sandia Mountains.
Discussions center on environmental concerns related tosurficial deposits and hydrology of the Llano de Sandia.
The route between Stops A-1 and A-2 traverses the Llanode Sandia between the mountain front and the Rio Grande
floodplain, via Tramway and Paseo del Norte Blvds. Stop A2 is at an exposure of ancient river deposits along Edith
Blvd. at the eastern edge of the floodplain. Discussion cen

ters on fluvial deposits of Pleistocene age that have been
dated by volcanic-ash correlation and vertebrate faunas.
We then cross the Los Ranchos de Albuquerque sectionof the floodplain on the new Paseo del Norte Freeway,
exiting on north Coors Blvd. for a tour of the western valley
border and Llano de Albuquerque area. This route follows
Irving Blvd. and Golf Course Rd. northward through Paradise Hills and Rio Rancho. We then turn west on Southern
Blvd. and traverse the dissected Llano de Albuquerque tableland west of Rio Rancho and northwest of the AlbuquerqueVolcanoes. Stop A-3 on the rim of the Ceja del Rio escarpment overlooks the Rio Puerco valley and fault zone,
which forms the western Rio Grande rift boundary. This isan excellent vantage point for viewing the eastern Colorado
Plateau, which extends from Mount Taylor to the Nacimiento Mountains, as well as the northern Albuquerquebasin and southern Jemez Mountains. Concepts of Santa Fe
Group stratigraphy and late Cenozoic landscape evolution,
originally developed in the northern Albuquerque basin by
Kirk Bryan (1909, 1938) and his students, are emphasized.
Following Stop A-3 we retrace our route east across the
Llano de Albuquerque on Southern Blvd. to the New Mexico
Gas Company pipeline road (northwestern extension ofAtrisco Dr.). The tour continues southeast on this road to Stop
A4 at the (Albuquerque) Volcano Cliffs overlooking Petroglyph Park (new National Monument area) and the inner Rio
Grande valley. Pleistocene activity of the AlbuquerqueVolcanoes and early history of valley cut and fill cycles are
discussed. The last segment of the route, via Atrisco Dr. andCoors Blvd., recrosses the valley floor on 1-40 and ends at the
1-25 south exit.
We recommend that this road log be used in conjunction
with V. C. Kelley's (1982) Scenic Trips to the Geologic Past
guidebook to the Albuquerque area, and the New Mexico
Geological Society Albuquerque Country II guidebook
edited by Grambling and Wells (1982). Lambert (1968),
Kelley (1977), Kelley and Northrop (1975), Kelley andKudo (1978), Machette (1978b, 1985), Clary et al. (1984),
Anderholm (1987, 1988), and Lozinsky (1988) should be

consulted for additional details on the environmental geologyof the Albuquerque—Rio Rancho area.
Directions and comments en-route to Stop A-1
Traffic light at Tramway Blvd. and Central Ave. (0.0 mi).
Travel north on Tramway Blvd. under 1-40 and across theupper piedmont deposits of Llano de Sandia (Fig. A-2). 1.4
mi, traffic light, crossing Lomas Blvd.; route ascends south
slope of Embudo Canyon fan. 2.0 mi, approaching intersection with Indian School Rd. Route crosses high axial partof fan; mouth of Embudo (funnel) Canyon in Sandia granite
at 3:00. 2.5 mi, descending north slope of fan; routebeneath the river floodplain 6 mi (10 km) to the west
(Kelley, 1982, pp. 41-43). The essentially flat water tablefrom here to the Rio Grande is indicative of very permeable
and thick basin fill (Kaehler, 1990; Kernodle and Scott,1986). Fig. A-3 shows the profiles of the land surface andwater table across this part of the Albuquerque basin.
4.7 mi, crossing Embudito Arroyo; note large granite
boulders. 5.1 mi, crossing main Bear Canyon Arroyo; floodcontrol structure downstream at 8:30. Granite boulders arecommon in coalescent-fan deposits here, but these boulders
decrease in size and abundance northward due to increasing
distance from the mountain front and flattening of the piedmont-slope profile. 5.5 mi, traffic light, crossing Spain Rd.
and northern distributary channel from Bear Canyon. 6.0
mi, traffic light, crossing Academy Blvd.; prepare for rightturn. 6.3 mi, turn right (east) on road to Elena Gallegos
Municipal Park, continue up valley of Arroyo del Pino. 7.8
mi, Elena Gallegos Park entrance booth. Stop A-1 parking
area is 0.4 mi to the east.

crosses main distributary channel from Embudo and Embudito fans. Small fan spreads out from mouth of Piedra
Lisa (smooth rock) canyon at 2:30. 4.0 mi, traffic light,
crossing Comanche Rd.; ascend south slope of Embudito
Canyon fan.
4.5 mi, traffic light at Montgomery Blvd. Road narrows;
continue north on Tramway Blvd. Ponderosa Reservoir andWell are 1.25 mi downslope at an elevation of about 5675 ft(1730 m). The 1242 ft (378 m) well is developed in Santa Fe
Group basin fill, and it is only 1.75 mi (2.8 km) from the
Sandia frontal fault mapped by Kelley and Northrop (1975).
Static-water-level elevation is about 4950 ft (1510 m),
approximately the same elevation as the water table

Stop A-1: Elena Gallegos Park and Sandia Mountainpiedmont. This stop (elevation 6500 ft, 1980 m) is in alarge embayment in the Sandia Mountain front known as
the Bear/Pino reentrant (Fig. A-2). Kelley (1982, p. 72;
Kelley and Northrop, 1975) suggested that the embayment
is caused by the eastward deflection of the Sandia fault
opposite the middle part of the mountains. He also considered that this part of the Sandias was uplifted first andsubjected to erosion longer than adjacent areas. Drainage
from three large canyons enters the embayment: Bear(south), Pino (central), and Domingo Baca (north). The
most prominent alluvial fan is at the mouth of Cañon
Domingo Baca. The material making up this fan is typicalof the older-piedmont sediment near the mountain front. It
consists of arkosic sandy gravel containing numerous
granite boulders 2.5 m in diameter, with some as large as
3.5 x 11 m. On the main part of the Llano de Sandia (acoalescent-fan piedmont or bajada extending about 8 mi, 13km, to the west) arkosic sandy pebble gravel to pebbly sandis the dominant lithology.
Presumably an extensive buried pediment, or suballuvial
bench, cut on Sandia granite occurs at shallow depths eastof the buried trace of the Sandia fault (about 0.5 mi, 0.8km, east of Tramway Blvd.; Kelley and Northrop, 1975, p.

13778). Within the Bear/Pino reentrant, granite is exposed insome of the deeper arroyo bottoms for distances as much as
1 mi (1.6 km) out from the mountain front (Lambert, 1968).
Pediments with well-developed rock-fan morphology(Johnson, 1932) and discontinuous alluvial cover are present
on the uppermost piedmont slope between the alluvial fan of
Cation Domingo Baca and Rincon Ridge (northwesternridge of the Sandias). Gravity studies by the U.S. Geological
Survey (Cordell, 1978a, b) indicate that a buried, major
boundary fault of Sandia uplift may be present about 3 km
(2 mi) west of this stop.
The undulating, upper-piedmont terrain at the southernbase of the Sandias is characterized by alternating alluvial
fans and interfan areas (Kelley, 1982, pp. 27, 28). Fans at
the mouths of larger canyons form major topographic highs
and extend about 3 mi (5 km) from the mountain front.
Broad topographic lows between the proximal parts of largefans commonly receive drainage from fan distributaries aswell as from adjacent mountain fronts and flanking alluvial
aprons. The latter include small fans and cones derived fromsmall, steep canyons. The large Bear (Oso), Embudo, andEmbudito fans to the south coalesce downslope to form a
relatively smooth landscape that grades to valley-border surfaces along the inner Rio Grande valley.
Whereas much of the Pleistocene runoff spread out on the
piedmont slope in fan distributaries, more recent concentration of flow in arroyos has resulted in landscape dissection. With gradients ranging from 3 to 6%, the present
arroyo system (and late Quaternary precursors) has produced
high-energy flows of water and sediment with competence tomove clasts up to large-boulder size. The detention dam,
located just west of Tramway Blvd. on Arroyo del Pino, as
well as the numerous other flood-control structures observedon this tour are operated by the Albuquerque Metropolitan
Arroyo Flood Control Authority (AMAFCA). AMAFCA is alocal-government subdivision which was created in 1963 andhas a district covering the City of Albuquerque and portionsof the surrounding Bernalillo County. AMAFCA derives
revenues through property taxes to construct and maintain
flood-control structures. In the first 10 years of AMAFCA's
history, efforts were devoted to the construction of the Northand South diversion channels that cost $28,619,000, of which
the taxpayers contributed a total of $11,410,000. Since 1973,
several additional flood-control structures have been
completed in the Northeast Heights of Albuquerque and on
the West Mesa; these include Embudo Arroyo Canal andCalabacillas Arroyo structures.
The Juan Tabo Dam, located about 3 mi (5 km) to thesouthwest on Bear Canyon Arroyo, is designed for the following conditions: (1) flow from the probable maximum
precipitation (PMP; 6-hour PMP effective rainfall = 39.9 cm)
will not overtop the structure; (2) the crest of the emergency
spillway is at an elevation that will not be overtopped by the
100-year frequency flood; (3) a continuously open culvert,
primary spillway 2.4 m in diameter, provides a maximum
discharge of 54.6 m3/s (1930 ft3/s) during the 100-year
frequency flood; and (4) the outlier culvert will drain the dam
following a flood that uses all available storage within eight
hours from beginning of storm event. The dam is constructedfrom piedmont deposits excavated within the reservoir area.
The structure has a maximum height of 60 ft (18.1 m),
maximum length of 1820 ft (554.7 m), and a maximum width
at base of 280 ft (85.3 m). The slope of the upstream face anddownstream face are 3:1 and 2:1(horizontal to vertical), respectively. The elevation at the
crest of the dam is 5779.5 ft (1761.6 m), and the detention
capability at the crest of the emergency spillway is 564,717
m3 (458 acre-ft). The drainage area above the Juan Tabodam is 26.33 km2 (10.17 mi2). The basin heads at the crestof the Sandia Mountains (elevation —9400-9600 ft, 28652926 m) and includes both Bear Canyon and Embudito
Arroyos. The total relief between the crest of the mountains
and the primary spillway of the dam is 1235 m (4053 ft).
In contrast to areas to the south near Tijeras Canyon,
piedmont aggradation over most of the Llano de Sandia
continued through late Pleistocene time, and erosion surfaces on upper Santa Fe Group deposits have been buried byyounger alluvium. The latter materials are predominantly
arkosic sandy gravel to gravelly sand with loamy to silty
zones. As noted by Kelley (1982, pp. 70-73), these deposits
are commonly bouldery in a zone extending several miles
out from the mountain front. Surface sediments in most
topographically high areas on the piedmont slopes comprisea complex of stream, hyperconcentrated floodflow, and debris-flow deposits emplaced before entrenchment of the latest Pleistocene and Holocene arroyo-valley systems. Distal
piedmont facies interfinger with, and partly overlap, theriver-terrace deposits adjacent to the inner valley of the Rio
Grande.
Stable surfaces on piedmont deposits are characterized
by well-developed soil profiles with pedogenic horizons ofclay accumulation over weak secondary carbonate zones.
Such features typify the most extensive soils on the Llanode Sandia (Hacker, 1977). The amount and depth of pedogenic clay and carbonate translocation in these soils (aswell as other weathering features) are indicative of wettersoil-moisture regimes that have not occurred in this region
since Wisconsin glacial-pluvial times (>10,000 years B .P.).
The late Pleistocene fan surfaces are cut by numerous small
draws and several broad arroyo valleys with Holocene fills.
Soil development is weak or essentially absent in thesematerials except for increasing amounts of organic matterin surficial horizons as one approaches the mountain front.
Holocene valley fills on upper-piedmont slopes also comprise bouldery stream to debris-flow deposits and locallycontain artifacts and charcoal.
In the areas southwest and northwest of this stop, Holocene arroyos have locally built low fans and alluvial aprons
on the upper piedmont slope (e.g. Chelwood Park area between Lomas Blvd. and Indian School Rd., east of Juan Tabo
Blvd.). Engineering interpretations by the Soil Conservation
Service of soils developed in Holocene and upper
Pleistocene alluvium indicate that many surficial deposits
have limitations due to compressibility, susceptibility to piping, and consolidation under load, and pose significant geotechnical problems to engineering geologists and civilengineers. Underconsolidated piedmont deposits with highvoid ratios are subject to hydrocompaction when wetted andare thus particularly troublesome materials in terms of current and potential construction hazards (Clary et al., 1984).
Directions and comments en-route to Stop A-2
Return to Tramway Blvd. (0.0 mi) and continue northacross Arroyo del Pino. AMAFCA flood-control structure toleft. 0.5 mi, traffic light, crossing San Rafael Ave. Base of
Sandia Peak Tramway at 1:00 is between La Cueva rock-fan
area (12:30) and bouldery alluvial fan of Cation Domingo
Baca (2:00). Juan Tabo reentrant in the Sandia front

138
at 12:00 separates the main part of the uplift from Rincon
Ridge on the west. The ridge is composed of Precambrianmetamorphic rocks intruded by pegmatite dikes. 1.0 mi,
crossing paved flood-runoff channel at San Bernadino Dr.;
prepare for left turn. 1.2 mi, turn left on Paseo del NorteDr.
For the next 6 mi, route descends the Llano de Sandia
piedmont plain, dropping from an elevation of 6000 ft to5000 ft (1830-1525 m) at Stop A-2 on the edge of thefloodplain. The piedmont slope in this area is a coalescentfan surface dissected by subparallel, west-flowing streams
into a subdued ridge and swale topography. Arroyo
gradients are about 4%. Upper Pleistocene alluvium
underlies the low ridges and Holocene alluvium underlies
the grass- and shrub-covered surfaces in the arroyo bottoms.
The alluvium is predominantly arkosic sandy gravel andgravelly sand with loamy (sand-silt-clay) interbeds. 2.6 mi,flood-water and sediment-control structure to right.
3.8 mi, traffic light, crossing Ventura St. An east-west
gravity profile (Cordell, 1978a, b) made along Tramway
Rd. about 2 mi (3.2 km) north of Paseo del Norte shows asteep westward gradient with a north-south linear trend.
The belt of steep gravity gradient crosses the tour route inthis area and indicates the presence of a large buried fault(downthrown to the west). 4.8 mi, traffic light, crossing
Wyoming Blvd. In 1940, the 5025 ft (1531 m) Norrins
Realty Co. exploratory well no. 2 was drilled about 0.3 mi(0.5 km) southwest of this intersection (surface elevation
about 5350 ft, 1630 m). According to the driller's log, theupper 2150 ft (655 m) is dominated by coarse, granitederived basin fill interpreted by Stearns (1953, pp. 475476)
as "Santa Fe Formation." It should be noted that the upper
150 ft (46 m) of fill is uncemented and bouldery, and may
represent post-Santa Fe, middle to upper Pleistocene
alluvium. The lower 2874 ft (876 m) penetrated by the
Norrins well includes mainly "gray sand" and "gray shale"that possibly correlate with the lower part of the Santa Fe
Group (Zia Formation) and/or the Espinaso Formation (Oligocene).
5.2 mi, Federal Aviation Administration facility to left,
crossing Louisiana Blvd. ahead at light. 6.1 mi, traffic light,
crossing 1-25 and frontage roads. 6.4 mi, continue west on
Paseo del Norte. Journal Center to left. 6.7 mi, traffic light,
crossing Jefferson St. Route ahead is on distal part of Llanode Sandia. Alluvium is late Quaternary and is part of a local
fan apron that spreads out from the mouths of arroyo valleys
heading in the Bear/Pino reentrant area. 7.1 mi, railroad
crossing; Reeves Power Plant to left. This is the zone ofmergence of the Llano de Sandia piedmont slope with terrace
surfaces bordering the inner valley of the Rio Grande. Active
and abandoned gravel pits to the right and left are developedin Edith alluvium, which is the major source of gravel andsand in the metropolitan area. 7.4 mi, cross North Drainage
Canal, one of two major outlet channels for flood runoff in
Albuquerque. This canal system intercepts runoff from theentire Llano de Sandia area between 1-25 and the Tramway.
For the period of record (1968-1979), the maximum discharge
measured at the gaging station 1.7 mi north of this point was
142 m3/s (5000 cfs) on July 26, 1971 (U.S. Army Corps ofEngineers, 1979). 7.7 mi, gravelly Edith alluvium exposed
under power poles to left.
7.9 mi, Edith Blvd. overpass (Stop A-2) 0.3 mi to north.
8.1 mi, crossing mainline Santa Fe Railroad. Prepare forright turn at next traffic light. 8.4 mi, turn right (north) on

2nd St.; continue 0.2 mi to Ortega Rd. and then turn right(east) 0.5 mi on Ortega to Edith Blvd. 9.1 mi, turn left, go
north 0.1 mi, and park at east edge of Edith at roadcut
exposure of basin-fill and terrace deposits (elevation 5000
ft, 1525 m).
Stop A-2: North Edith section of Lambert (1968) andenvironmental geology of the eastern valley-border area.
An axial-river facies of the upper Santa Fe Group (Upper
Buff[?] sand of Lambert, 1968) is exposed in cut to right. In
the Albuquerque area, similar ancestral-river deposits ofearly Pleistocene age are discontinuously exposed in thelower part of the eastern valley-border scarp. In this area, thisfluvial unit has a maximum exposed thickness of 3343 ft (1013 m) and is overlain with erosional unconformity by
gravelly Edith alluvium. However, because of inadequate
exposures, the unit cannot be physically traced to well-dated
Sierra Ladrones Formation (Ceja Member) deposits that cropout along Tijeras Arroyo and the Llano de Albuquerqueescarpments to the south and west (Stops A-3 and A-4).
The axial-river facies is here composed mainly of waterdeposited rhyolitic pumice in the sand- and pebble-size
range. About 2 mi (3.2 km) north of this stop the waterdeposited pumice is conformably underlain by an air-fall ash
layer resting on clay to clayey sand (Lambert, 1976). The airfall ash and, presumably, the overlying water-deposited
pumice have long been regarded as derived from one or morecaldera-forming eruptions in the Jemez Mountains (Lambert,
1968, Appendix A; Smith et al., 1970). These eruptions
produced the two members of the Bandelier Tuff and theintra-Bandelier Cerro Toledo Rhyolite; all units are between
1 and 1.5 m.y. old (Bailey and Smith, 1978; Goff et al.,1989). A sample of the air-fall ash collected by Ray Wilcox(U.S. Geological Survey) and Lambert has been correlatedby Izett et al. (1981) with the Guaje Pumice Bed, the basal
unit of the Bandelier Tuff. The K-Ar age of the Guaje
Pumice is about 1.5 m.y., thus placing this part of the upper
Santa Fe Group in the early Pleistocene. This age fits wellwith other deposits of the ancestral Rio Grande (Sierra Ladrones, Palomas, and Camp Rice Formations) dated both by
K-Ar and zircon fission-track methods and by vertebrate
faunas (Hawley, 1978, chart 1; Gile et al., 1981; Tedford,
1981; Lozinsky and Hawley, 1986).
The overlying Edith alluvium is middle to upper Pleistocene Rio Grande terrace deposit with an approximate
thickness of 40-50 ft (12-15 m); however, at this stop theupper part has been removed by erosion and excavation. It is
composed of gray sandy gravel and light-colored overbank
mud and sand. The gravel is characterized by abundant,
well-rounded pebbles and cobbles of metaquartzite. The
Edith is exposed along the east side of the Rio Grande
between Albuquerque and Bernalillo and occurs on bothsides of the river in Albuquerque. It is locally inset against
older valley fill and contains middle to late Pleistocene(Rancholabrean) fauna (Lambert, 1968, pp. 174-178;
O'Neill and Rigby, 1982).
It has long been common practice to "reclaim" gravelpits in the Edith and other fluvial units with various typesof municipal and industrial wastes. Land-fill operations using abandoned gravel pits extend from here to the valley ofTijeras Arroyo. The danger of this practice, particularly interms of potential for ground-water pollution, is readily
apparent at this stop. The water table here is within 15 to 30
ft (5 to 10 m) of the surface, and the permeable Edithdeposits are in direct contact with the equally permeable
upper Santa Fe Group beds of unknown but significant
thickness and wide subsurface extent. It is quite probablethat the older fluvial beds seen here continue southeastward
under the Llano de Sandia, as well as other areas where theupper Santa Fe Group is a major component of the basinfill aquifer system.
Most cases of ground-water pollution are located near the
inner Rio Grande valley where depth to ground water isshallow (Wilson, 1981; Stone, 1984; Anderholm, 1987).
Regional, low-level pollution of the Rio Grande valley-fillaquifer with fecal coliform appears to be related to local
unsanitary conditions of on-site waste disposal systems
rather than to regional microbial pollution. The majority ofthese systems are septic tanks. In addition to pollution bynitrates, which has a 20-year record for portions of the Rio
Grande valley, pollution by organic compounds has been
detected by the Environmental Improvement Division (EID)of New Mexico. The primary pollutant is hydrocarbon-fueloils such as gasoline and diesel fuel. These pollutants havebeen found in a variety of locations in the Rio Grande valleyof the Albuquerque area. Federally funded grants to the State
EID are supporting studies of pollution by nitrates, hydrocarbons, and other toxic organics in the Rio Grande valley(McQuillan, 1982; Thomson and McQuillan, 1984).
Return to Paseo del Norte Blvd. via Ortega Rd. and 2nd
St.
Directions and comments en-route to Stop A-3
This road-log segment starts at the 2nd St. entrance tothe Paseo del Norte Freeway (westbound lanes-0.0 mi). For
the next 2.4 mi, the freeway is on the Rio Grande
floodplain, crossing several irrigation canals, drains, andabandoned channel segments, as well as the modern river
channel and floodway (1.5 mi). Floodplain gradient is 5 ft/
mi (<1 m/km), which contrasts markedly with the greaterthan 50 ft/mi (>10 m/km) slopes of valley-border and piedmont terrains. At Albuquerque, upper Quaternary river deposits, dating from the past 20,000 years, are about 75 ft(23 m) thick according to Kelley (1982, pp. 13, 90-93).
They include a basal gravel and sand zone that grades upward to sand with gravel lenses and thin surficial layers ofclay and silt. This fluvial-facies assemblage, locally as
much as 130 ft (40 m) thick, characterizes Rio Grandevalley-floor deposits from the Albuquerque basin to Huecobolson below El Paso, Texas (Hawley et al., 1976; Gile etal., 1981). Basal coarse-grained deposits relate to episodesof valley cutting during late Wisconsin glacial maxima.
The upper sand-to-clay zone represents Holocene and latest
Pleistocene aggradation, a process characteristic of stream
valleys throughout the region.
This part of the city is susceptible to flooding caused byrainstorm-runoff events in local valley-border (Fig. A-4) andpiedmont watersheds rather than by overbank flood discharges of the river. Large river floods are now controlled by
upstream structures such as Cochiti Dam (Lagasse, 1981).
Historical peak flows of the middle Rio Grande (1866-1967;
as much as 100,000 cfs, 2850 m3/s) are shown in Fig. A5(Kelley, 1982, p. 19). Stay in right lane for Coors Blvd. exit
at west end of Paseo del Norte Freeway.
2.3 mi, take Coors Blvd. exit; keep to right. 2.5 mi, takeNorth Coors Blvd. exit. Continue 0.5 mi north on Coors to
Irving Blvd.; quickly move to far left lane for left turn. 3.0
mi, traffic light. Turn left (west) on Irving Blvd. Route rises

across dissected valley-border surfaces (fans and terraces)of middle to late Quaternary age.
3.8 mi, crossing terrace overlooking Arroyo de las Calabacillas (to right). This arroyo system drains much of thenorthern Llano de Albuquerque (basin area of about 220
km2). Leopold (1946) described the effects of a severe thunderstorm that moved northeastward across the Llano de Albuquerque northwest of Albuquerque on the afternoon of
September 20, 1941. Rainfall amounts of as much as 3864
mm may have been produced by the storm. Peak flow at themouth of Arroyo de las Calabacillas was estimated to havebeen 285 m3/s (10,000 ft3/s). Arroyos north of Arroyo de las
Calabacillas had estimated peak flows of 60-115 m3/ s(2000 to 4000 ft3/s). According to local testimony, a similar
flash flood occurred in Arroyo de las Calabacillas in 1904.
The fan deposited at the mouth of Arroyo de las
Calabacillas by the 1941 flood was described by Leopold(1946, p. 539) as follows:
At the mouth of Calabacillas Arroyo, an estimated 250 acre-ft
of sediment, primarily sand, were deposited as a fan which
forced the flow of the Rio Grande into a narrow channel
against the east bank of the main floodplain. The total contribution of sediment must have been even larger becausemuch was carried downstream by the combined flow of thearroyo and the Rio Grande. The depth of deposition was ninefeet in places. Some large boulders were deposited in the fan.
One, estimated weight 1,400 pounds, was deposited 700

140
Blvd. Route ahead crosses broad undulating surface on upper Santa Fe Group veneered with Quaternary alluvium andeolian deposits. 8.7 mi, traffic light, crossing 20th St. 9.4
In 1941, the Arroyo de las Calabacillas area was largely
mi, route ascends slope cut in Ceja Member of Sierra Laundeveloped. The only damage reported by Leopold was
drones Formation gravelly sand. This is the eroded scarp
the washing out of a primary gas main feeding the City of
Albuquerque. If a similar flood occurred today, damage of Kelley's (1977) Star Heights fault.
10.3 mi, on crest of high ridge about 650 ft (200 m)
could be much more severe (Fig. A-4).
above
floodplain level. Mount Taylor on skyline. For next
4.6 mi, traffic light. Turn right (north) on Golf Course Rd.
8.3 mi route crosses upper Arroyo de las Calabacillas wateracross channel of Arroyo de las Calabacillas. Construction offlood-water-retention structures in this area has been plagued shed (Fig. A-6). An outlier of the Llano de Albuquerqueby problems because of susceptibility of sandy-substrate units north of this point is locally capped with a strong calcic soil
to piping and seepage, and lack of nearby sources for suitable developed on pebbly sand of the Ceja Member, Sierra Ladrones Formation. The water table here is about 700 ft (213
clayey fill materials. 5.0 mi, traffic light at Ellison. 5.7 mi,m) below the surface and has a trough-like configuration
crossing Black Arroyo. 5.9 mi, Bernalillo—Sandoval County(Fig. A-1), sloping southward away from adjacent parts ofline; entering the City of Rio Rancho (New Mexico's fastest
the Rio Grande and Rio Puerco valley floors (Titus, 1961;
growing urban area). Continue north on Golf Course Rd.
Kernodle and Scott, 1986; Anderholm, 1987, 1988; Hawley
across dissected, older valley-border surface (about 246-263
and Longmire, in press). Wells for public water supplies in
ft, 75-80 m, above the valley floor) that is part of the midthis part of the Albuquerque basin are commonly more thanPleistocene Tercero Alto geomorphic-surface complex of1000 ft (300 m) deep and are developed in the middle toMachette (1985). The surface is locally covered by basalts ofupper part of the Santa Fe Group. Near the western rim of
the Albuquerque Volcanoes field (Stop 4) dated at <200,000
the Llano de Albuquerque, depths to the basin-fill aquifer
yrs by Geissman et al. (1990). Middle to late Quaternary
locally exceed 1000 ft (300 m).
alluvial, colluvial, and eolian deposits mantle Plio-Pleistocene
11.5 mi, on graded road descending into valley of Arroyo
basin fill of the upper Santa Fe Group—Sierra Ladronesde las Calabacillas across slopes cut in sandy basin fill (upper
Formation (Ceja MemSanta Fe Group). 11.7 mi, caution, dip in road ahead atber?).
Arroyo de las Calabacillas. Southern Blvd. occasionally
7.3 mi, traffic light ahead, prepare for left turn. 7.5 mi,
washes out at this point during large storm-runoff events. For
turn left (west) on Southern Blvd. 7.7 mi, Albuquerquenext 3.5 mi route is on thin veneer of upper Quaternary
Volcanoes at 9:30. Continue west about 10 mi on Southernfeet from the mouth of the arroyo and five feet above thenormal streambed.

valley fill underlain by upper Santa Fe Group basin fill. 12.9
mi, on line with north—south trend of Albuquerque Volcanoes (Kelley and Kudo, 1978; Geissman et al., 1990). The
northernmost cone is about 5 mi (8 km) to the south ateastern edge of the Llano de Albuquerque surface. In thisarea the route crosses at least two faults that are along thetrend of the Albuquerque Volcanoes fissure. Displacement isdown to the east. These faults line up to the south with theCounty Dump—Nine Mile fault trend of Machette (1978a)and Kelley (1977), and they are aligned to the north with theDuval and Zia faults of Kelley (1977). Machette's (1978c)study of the County Dump fault, near 1-40, shows that there
has been recurrent movement on this feature during middle
to latest Pleistocene time.
13.4 mi, intersection with Gas Company of New Mexico
pipeline road. (Route to Stop A-4.) 15.2 mi, near top of scarpthat forms the western edge of the Arroyo de las Calabacillas
valley. Coarse-grained fluvial deposits of the Ceja Member,
Sierra Ladrones Formation, form the upper unit of the Santa
Fe Group in this area. Clasts are in the pebble-to-smallcobble range, commonly rounded, and include quartzite,
chert, basalt, and granite. Some of the chert is derived fromthe Pedernal Chert Member of the Abiquiu Formation(Oligocene—Miocene; see Vazzana and Ingersoll, 1981). A
calcic soil is developed in the uppermost sand and gravel
beds and forms the light-colored (Cejita Blanca) rim of the
Llano de Albuquerque ahead. This strong horizon of soilcarbonate accumulation is probably correlative with calcic
horizons in surficial soils of the eastern Llano de Albuquerque scarp seen at Stops 2-2 and 2-4 on Day 2.
15.4 mi, military training area to right; crossing remnantof Llano de Albuquerque. This early to middle Pleistocene
geomorphic surface represents a relict central-basin plain that
predates entrenchment of the present Rio Grande and Rio
Puerco valleys (Lambert, 1968; Machette, 1985). The
original basin-floor and distal-piedmont components of thisconstructional landscape have been faulted, partly dissectedby erosion, and locally buried by later eolian and alluvialcolluvial deposits. The Llano surface to the south and west is
crossed by north-trending linear depressions at the base oferoded fault scarps. Offsets are mainly down to the east
(Kelley, 1977; Machette, 1982). 16.5 mi, crossing broad,north-trending swale. 17.0 mi, note high dunes on the skyline
extending south from the AT&T Co. relay tower (10:001:00). These dunes partly cover the western rim of the Llanode Albuquerque and are derived from the Rio Puerco valley
area to the west (Lambert, 1968, 1974).
17.8 mi, turn right (north) on 60th St. 18.0 mi, opposite
AT&T Co. relay installation. Turn left (west) on dirt road.
18.1 mi, powerline crossing. Road ahead jogs to right andthen to left. Route ascends Holocene alluvial and eolian
apron at eastern base of high rim dunes. 18.4 mi, keep left
at road fork. Continue west to base of large dune and parkfor Stop A-3 (18.6 mi). Road to right is recommended only
for offroad vehicles due to loose sand. It skirts the northernend of the rim dunes and leads to the Ceja del Rio Puerco([eye]brow of the Rio Puerco [escarpment]).
Stop A-3: Ceja del Rio Puerco and Sand Hill fault zone.
Walk to crest of dune for orientation and then west to Sand
Hill fault zone at rim of Ceja del Rio Puerco escarpment(Figs. A-1, A-6, A-7).
The rim-dune crest (elevation 6263 ft, 1909 m) offers abroad vista of the northern Albuquerque basin and adjacentparts of the Colorado Plateau and Southern Rocky Mountainprovinces. Mount Taylor (west) and the Lucero uplift(south-southwest) flank the Acoma sag area of the Colorado
Plateau visited on Day 2 (Stop 2-1). Basalt-capped Mesa
Prieta, about 12 mi (20 km) to the northwest, is located eastof the Rio Puerco valley (see Slavin, this volume). The
Nacimiento Mountains to the north and the Sangre de Cristo
Mountains on the northeastern skyline form thesouthernmost prongs of the Southern Rockies. The Jemez
Mountain volcanic center (north-northeast) of middle
Miocene to Pleistocene age is between these ranges (Bailey
and Smith, 1978; Goff et al. , 1989).
The base of the escarpment in the foreground is along
the eastern margin of the Rio Puerco fault zone (Kelley andWood, 1946; Wright, 1946; Slack and Campbell, 1976).
About 4 mi (6.4 km) west of here the lower part of the
Santa Fe Group is down-faulted against Upper Cretaceous
rocks along the Moquino fault of Kelley (1977; Canada
Moquino fault of Bryan and McCann, 1937). This fault andthe Sand Hill fault, which crosses the Ceja rim at this stop,
appear to form the two major rift-boundary faults in thisarea (Bryan and McCann, 1937; Wright, 1946; Kelley etal., 1976; Kelley, 1977; Machette, 1982).
This stop is near the southern end of the Ceja del Rio
Puerco area where Bryan and McCann (1937) first subdivided the "Santa Fe formation" into broad mapping units(Tedford, 1982). Their threefold subdivision of the Santa
Fe into "Lower Gray, Middle Red, and Upper Buff members" was based primarily on variations in color, texture,

142
and cementation. They characterized the Lower Gray andMiddle Red members as having mainly sandy textures, withclay and gravel beds and calcareous concretionary sandstones locally present. The main distinguishing feature wascolor, white to gray in the Lower Gray unit and common
shades of red to buff in the Middle Red unit. The presenceof gravel, buff colors, and variable textures characterized
their Upper Buff member. These basic Santa Fe
subdivisions have been used (with some modifications) bymost subsequent workers in the northern Albuquerquebasin, including Wright (1946), Campbell (1967), Lambert(1968), and Kelley (1977); they are still commonly used forgeneral descriptions of basin-fill stratigraphy. Even thesebroad Santa Fe mapping units have been useful in making
estimates of minimum throw on major basin-margin andintrarift faults, and they allowed Bryan and McCann (1937,
p. 801) "to speculate on the amount of the depression of thebasin as a result of this faulting."
Galusha (1966) mapped the Lower Gray and Middle Red
type area in the Rincones de Zia about 20 mi (33 km) to thenorth, and he made the first definitive studies of vertebrate
faunas and biostratigraphy. Galusha formally named thelower to middle Miocene Zia Sand Formation (now called
the Zia Formation), and showed that it was equivalent to
the type Lower Gray member of Bryan and McCann(1937). Galusha and Blick (1971, pp. 39-40, 70) demonstrated that at least the lower part of the Middle Red
member correlated with the upper part of the middle
Miocene Tesuque Formation, the main type Santa Fe
subdivision in the Española basin (Spiegel and Baldwin,1963). Gawne (1973, 1981) refined the Zia Formation
stratigraphic nomenclature in the Canada Pilares—Arroyo
Benavidez area (2.5 mi, 4 km) to the north-northwest of
this stop, as well as in the type Zia area along the valley of
the Jemez River. Kelley (1977) described the Zia and lower
Middle Red units at the Benavidez Ranch measured section
about 3.5 mi (5.6 km) west-southwest of this stop.
Current work by Tedford and Barghoon of the American
Museum of Natural History involves completion of biostratigraphic studies in the northern Ceja area that were
initiated by the late Ted Galusha. Detailed measured
sections have been made between here and Canada Pilares,
and work on magnetostratigraphy as well as biostratigraphy
is in progress. The exposed Santa Fe Group section is about1000 ft (300 m) thick and thickens southeastward. The
lower 415 ft (126 m) is the Zia (Sand) Formation describedby Gawne (1981). This predominantly eolian sand unit isangularly unconformable on Upper Cretaceous sandstoneand shale with very thin remnants of the Galisteo
Formation locally present between the Cretaceous and basal
Zia. The upper Zia, Canada Pilares Member of Gawne(1981), a red clay with local sandstone lenses up to 92 ft(28 m) thick, has been included with the Middle Red
member by previous workers. The upper part of the Middle
Red member of Kelley (1977) is here about 650 ft (200 m)
thick and consists mainly of light reddish-brown sand tosilty sand. The unit has thin beds of volcanic ash and one
prominent zone of algal tufa heads and lacustrine clay tosand. The "Middle Red member" contains eolian-sandbodies very similar to the bulk of the underlying Zia
Formation, and it is included in that formation by Tedford(pers. comm. 1990) as an unnamed upper member.
The upper Santa Fe Group at this stop includes the Ceja

Member of Kelley (1977), a sand-and-gravel to conglomeratic sandstone facies with local thick lenses of loamy toclayey sediments and buried soils in structural sags
adjacent to faults. The Ceja Member as mapped by Kelley(1977) includes much of the Upper Buff member of Bryan
and McCann (1937), Wright (1946), and Lambert (1968),
and it is a basin-floor facies of the Sierra Ladrones
Formation of Machette (1978a, b). Along the escarpment
near this stop, the Ceja Member is as much as 250 ft (76 m)
thick and the upper 50-100 ft (15-30 m) of the unit containscattered angular boulders of basalt and Cretaceous
sandstone, probably derived from the area to the northwestthat is now part of the Rio Puerco valley. The bulk of thegravel clasts are rounded pebbles of mixed siliceous rocktypes derived from as far north as the northern Nacimiento
uplift. The Ceja Member is capped with a strong calcic soil
and thick eolian sand of the rim-dune sequence.
Tedford (1981, 1982) reviewed the mammalian biochronology of the Santa Fe Group and the development ofstrati-graphic terminology in the Ceja del Rio Puerco andJemez Valley areas. The lower Zia Formation contains
faunas of late Arikareean and Hemingfordian provincial
age (about 21 to 16 Ma). Sites in the Middle Red member(upper Zia of Tedford) have late Barstovian and
Clarendonian faunas (about 15 to 10 Ma). Faunas of the
Ceja Member, Sierra Ladrones Formation, are primarily ofBlancan age (less than 4.5 Ma).
Nearby volcanics within the Rio Puerco valley include
the basalt flow capping Mesa Prieta about 12 mi (20 km) to
the northwest, numerous basaltic necks, and the Benavidez
diatreme 3 mi (5 km) west of this stop. The Mesa Prieta
flow with a K—Ar age of 2.2 ± 0.3 m. y. (Armstrong et al.,1976) rests on an erosion surface correlated with the Ortiz
surface by Bryan and McCann (1938). The Benavidez diatreme, first described by Kelley and Kudo (1978), is associated with a partly exhumed maar center emplaced in thelower and middle parts of the Santa Fe Group section. Volcanic necks and geomorphic evolution of the region to thenorthwest are discussed by Slavin (this volume).
The Sand Hill fault zone of Bryan and McCann (1937) is
well exposed at this stop (Fig. A-6). It was first described indetail by Wright (1946). He noted the episodic movement
and growth of the main fault segment exposed at this stop,
which involves offset (down to east) of both Middle Red
and Upper Buff members. As at the County Dump fault(Machette, 1978c), episodic fault displacement alternatingwith intervals of erosion of the upthrown block has resulted
in episodic sediment accumulation and soil development on
the downthrown block.
The lower Rio Puerco valley area, including summits offlanking tablelands (e.g. Llano de Albuquerque), contains
several proposed sites for large solid-waste management
facilities (Hawley and Longmire, in press). A regional landfill for Bernalillo County (Cerro Colorado facility) has already been developed on the Ceja del Rio Puerco about 15
mi (24 km) south of this stop. Proposed sites are characterized by thick vadose zones (500 to 1000 ft, 150-300 m) in
fine- to medium-grained Neogene basin fill or Cretaceous
rocks and absence of large fresh-water aquifers. Siting problems include (1) avoiding geomorphically unstable areas
with potential for rapid erosion—sedimentation rates or intense seismic activity, and (2) protecting ecologically andculturally sensitive areas.

143
Directions and comments en-route to Stop A-4
Retrace route 5.2 mi via Southern Blvd. (4.4 mi) to junction with Gas Company of New Mexico pipeline road (Figs.
2 and A-6). 0.0 mi, turn right (southeast) on pipeline road(graded extension of Atrisco Dr.—Unser Blvd.). Route descends erosion surface cut on upper Santa Fe Group—Sierra
Ladrones Formation (Ceja Member?). 4.0 mi, crossing basalt flows of Albuquerque Volcanoes field from here to Stop
A-4; note extensive eolian cover. 6.1 mi, start of pavementat Unser Blvd.; approaching Volcano Cliffs. Turn right(southwest) on paved road to Maloof Memorial Airpark and
La Boca Negra Park for 0.15 mi; turn left (southeast) at
power poles on gravel (pipeline) road and continue 0.45 mi
past pipeline enclosure; turn left and continue east to edgeof basalt flow and Stop A-4 (6.8 mi). Pipeline crosses rim ofVolcano Cliffs escarpment at this point (elevation 5300 ft,
1615 m).
Stop A-4: Albuquerque Volcanoes and Volcano Cliffs.
Kelley and Kudo (1977) mapped ten basalt flows in the
Albuquerque volcanic field, including an extensive basal (Qbl)
flow of olivine basalt capping cliffs at this top (Fig.

A-8). This flow has been K—Ar dated at 190,000 ± 40,000
years by Bachman et al. (1975). According to Machette(1978c, pp. 374-375), the most extensive flows spread easterly over the two highest Rio Grande terraces (Tercero and
Segundo Alto) between about 200,000 and 100,000 years
ago. Geissman et al. (1990) show that all flows of the
Albuquerque Volcanoes record a short-excursion, magneticpolarity episode. K—Ar age determinations (weighted average: 155 ± 47 ka) and morphology of calcic soils developed in flow veneers enriched in eolian materials suggestthat the short extrusion episode occurred sometime between
250 and 80 ka.
A series of at least ten small volcanoes and spatter cones
are aligned along the trend of the County Dump fault—
Albuquerque Volcanoes fissure (Lambert, 1968). The Vulcan "J" Cone center is the most prominent topographic
feature. The cone is about 200 ft (60 m) high and 2100 ft(640 m) in diameter. It is a steep-sided cinder and lava cone
flanked on the north by a questionable collapse structure
and on the south by two small adventive cones.
Poorly consolidated sediments, locally exposed beneath theflows, include pink and green clay and silt as well as clean,light-colored sand. These beds were originally mapped

144
by Lambert (1968, pp. 73, 103) as Upper Buff member (Ceja
Member equivalent) of the Santa Fe Formation. They may
instead correlate with older valley fill beneath the "Tercero
Alto" surface of Machette (1985), which includes terrace
alluvium in the Volcano Cliffs area that is associated withsurfaces about 240 ft (73 m) above the valley floor. An older,axial-river gravel that is poorly exposed in low hills at thebase of the Volcano Cliffs escarpment north and south of thisstop was mapped by Lambert (1968, p. 94) as Upper Buff(?)
gravel. This cobbly, erosion-resistant unit, which is up to 50
ft (15 m) thick, is probably an exhumed channel zone within
the upper (Pliocene—early Pleistocene) part of the Sierra
Ladrones Formation. The channel gravel crops out
discontinuously along the west border of the river valley
from near Bemalillo to Isleta Pueblo, where it is offset by atleast one major fault and may be partly buried by a basalt
flow of the Isleta volcanic center, with a K—Ar age of 2.78
Ma (Kelley and Kudo, 1978; Kudo et al., 1977).
Directions and comments en-route from Stop A-4
to I-40-1-25 Interchange
Retrace route to Unser Blvd. (0.0 mi). Continue southeast
on Unser, descending Volcano Cliffs escarpment. 0.5 mi,entrance to Indian Petroglyph State Park on left. 1.0 mi,intersection with Montaño Blvd. San Antonio Arroyo floodwater channel on right cut in exhumed cobbly river-channel
deposit in the Sierra Ladrones Formation, possibly as old aslate Pliocene. This is the Upper Buff(?) gravel of Lambert(1968). 1.8 mi, drainage channel and powerline crossing at
foot of basalt-capped spur of basin- and/or valley-fill alluvium, with mantle of basalt-block colluvium. Route ahead
on Atrisco Dr. descends to Segundo Alto constructional
terrace, here about 140 ft (43 m) above the Rio Grande
floodplain. The terrace fill comprises the Los Duranes alluvium (alloformation) of Lambert (1968) and locally exceeds 140 ft (43 m) in thickness. Deposits are primarily
sand, silt, and clay, with a basal gravel zone. One flow
tongue of the Albuquerque Volcanoes is exposed in a valley
cut below the Segundo Alto surface; it may have beenburied during deposition of the underlying Los Duranes
alluvium and subsequently exhumed.
2.9 mi, intersection of Atrisco Dr. and Ladera Dr.; turnleft (east) on Ladera and prepare for right turn. 3.2 mi,intersection with Coors Blvd. Turn right on Coors and continue 1.5 mi south across Segundo Alto terrace. Stay incenter lane; prepare to take second exit to right (1-40 East).
4.7 mi, take 1-40 East (Santa Rosa) exit to right. 4.9 mi,
merge with 1-40 and continue east across inner Rio
Grande valley to interchange with 1-25. Route descends
from Segundo Alto surface to river-valley floor. Roadcuts
ahead in Los Duranes alluvium. 5.6 mi, crossing Rio
Grande channel and floodway. Route on Rio Grande
floodplain for next 3 mi. Historic Old Town Plaza ahead
on right at Rio Grande Blvd. exit. 6.7 mi, Rio Grande
Blvd. exit to right; Old Town Plaza, AlbuquerqueMuseum, and New Mexico Museum of Natural History.
7.4 mi, 12th St. exit to right. Pueblo Museum three blocks
north of freeway on 12th. I40 and 1-25 ("Big I")
Interchange ahead. Prepare to take 1-25 South. 8.9 mi,take 1-25 South exit to University of New Mexico (Lomas
and Grand exits), and to Airport (Girard exit). End of
Albuquerque urban-area tour and Day 3 of field trip.

References
Akers, J. P., 1964, Geology and ground water in the central part of ApacheCounty, Arizona: U.S. Geological Survey, Water-Supply Paper 1771,
107 pp.
Anderholm, S. K., 1987, Reconnaissance of hydrology, land use, groundwater chemistry, and effects of land use on ground-water chemistry in
the Albuquerque—Belen basin, New Mexico: U.S. Geological Survey,
Water-Resources Investigations Report 86-4174, 37 pp.
Anderholm, S. K., 1988, Ground-water geochemistry of the AlbuquerqueBelen basin, central New Mexico: U.S. Geological Survey, WaterResources Investigation Report 86-4094, 110 pp.
Anderson, 0. J., Love, D., and Lucas, S. G., 1989, Supplemental road log
3, from intersection of NM-602 and Bread Springs Road to Fort
Wingate, via Stinking Springs and McGaffey: New Mexico GeologicalSociety, Guidebook 40, pp. 71-73.
Anderson, 0. J., Lucas, S. G., and Love, D. W., 1989, Stratigraphic
nomenclature chart: New Mexico Geological Society, Guidebook 40, p.
xiv.
Anderson, 0. J., Lucas, S. G., Love, D. W., Maxwell, C. H., andChamberlin, R. M., 1989, Third-day road log, from Gallup to Upper
Nutria, Ramah, El Morro and Grants: New Mexico Geological Society,
Guidebook 40, pp. 49-66.
Antevs, E., 1954, Climate of New Mexico during the last glacial—pluvial:
Journal of Geology, v. 62, pp. 182-191.
Armstrong, R. L., Speed, R. C., Graustein, W. C., and Young, A. Y., 1976,
K—Ar dates from Arizona, Montana, Nevada, [New Mexico], Utah, andWyoming: Isochron/West, no. 16, pp. 1-6.
Bachhuber, F. W., 1971, Paleolimnology of Lake Estancia and the Quaternary history of the Estancia Valley, central New Mexico: Unpublished
PhD dissertation, University of New Mexico, Albuquerque, 238 pp.
Bachhuber, F. W., 1982, Quaternary history of the Estancia Valley, central
New Mexico: New Mexico Geological Society, Guidebook 33, pp. 343346.
Bachhuber, F. W., 1989, The occurrence and paleolimnologic significanceof cutthroat trout (Oncorhynchus clarki) in pluvial lakes of the Estancia
Valley, central New Mexico: Geological Society of America, Bulletin,v. 101, pp. 1543-1551.
Bachman, G. 0., and Mehnert, H. H., 1978, New K—Ar dates and the late
Pliocene to Holocene geomorphic history of the central Rio Grande
region, New Mexico: Geological Society of America, Bulletin, v. 89,
pp. 283-292.
Bachman, G. 0., Marvin, R. F., Mehnert, H. H., and Merritt, V., 1975,
K—Ar dates of Los Lunas and Albuquerque basalts: Isochron/West, no.
13, pp. 3-4.
Bailey, R. A., and Smith, R. L., 1978, Volcanic geology of the Jemez
Mountains, New Mexico; in Hawley, J. W. (compiler), Guidebook to
Rio Grande rift in New Mexico and Colorado: New Mexico Bureau of
Mines & Mineral Resources, Circular 163, pp. 184-196.
Baldridge, W. S., Perry, F. V., and Shafiqullah, M., 1987, Late Cenozoic
volcanism of the southeastern Colorado Plateau, I. Volcanic geology of
the Lucero area, New Mexico: Geological Society of America, Bulletin,v. 99, pp. 463-470.
Baldridge, W. S., Perry, F. V., Vaniman, D. T., Nealey, L. D., Leavy,
D. B., Laughlin, A. W., Kyle, P., Bartov, Y., Steinitz, G., and Gladney,
E. S., 1989, Excursion 8A: Magmatism associated with lithosphericextension: Middle to late Cenozoic magmatism of the southeastern Colorado Plateau and central Rio Grande rift, New Mexico and Arizona; in
Chapin, C. E., and Zidek, J. (eds.), Field excursions to volcanic terranes
in the western United States, Volume I: Southern Rocky Mountain region: New Mexico Bureau of Mines & Mineral Resources, Memoir 46,
pp. 187-202.
Baldridge, W. S., Perry, F. V., Nealey, L. D., Laughlin, A. W., andWohletz, K. H., 1989, Field guide to Excursion 8A: Oligocene to Holocene magmatism and extensional tectonics, central Rio Grande rift andsoutheastern Colorado Plateau, New Mexico and Arizona; in Chapin,
C. E., and Zidek, J. (eds.), Field excursions to volcanic terranes in thewestern United States, Volume I: Southern Rocky Mountain region:
New Mexico Bureau of Mines & Mineral Resources, Memoir 46, pp.
202230.
Bates, R. L., Wilpolt, R. H., MacAlpin, A. J., and Vorbe, G., 1947,
Geology of the Gran Quivira quadrangle: New Mexico Bureau of Mines& Mineral Resources, Bulletin 26, 52 pp.
Betancourt, J. L., VanDevender, T. R., and Martin, P. S. (eds.) 1990,
Fossil packrat middens: The last 40,000 years of biotic change: University of Arizona Press, Tucson, 469 pp.
Blagbrough, J. W., 1865, Quaternary geology of the northern Chuska

145
Mountains and Red Rock Valley, northeastern Arizona and northwestern
New Mexico: Unpublished PhD dissertation, University of New Mexico,
Albuquerque, 138 pp.
Bourlier, B. G., Neher, R. E., Crezee, D. B., Bowman, K. J., and Meister,D. W., 1970, Soil survey of Torrance area, New Mexico: U.S. Soil
Conservation Service, U.S. Government Printing Office, Washington,
D.C., 149 pp.
Brackenridge, G. R., 1978, Evidence for a cold, dry full-glacial climate in
the American Southwest: Quaternary Research, v. 9, pp. 22-40.
Brady, B. T., Mulvihill, D. H., Hart, D. L., Jr., and Langer, W. H., 1984,
Maps showing ground-water levels, springs, and depth to groundwater, Basin and Range province, New Mexico: U.S. Geological
Survey, Water-Resources Investigation Report 83-4118-B.
Brod, R. C., and Stone, W. J., 1981, Hydrogeology of Ambrosia Lake-San
Mateo area, McKinley and Cibola Counties, New Mexico: New
Mexico Bureau of Mines & Mineral Resources, Hydrogeologic Sheet2, scale 1:62,500.
Brown, L. D., Chapin, C. E., Sanford, A. R., Kaufman, S., and Oliver, J.,
1980, Deep structure of the Rio Grande rift from seismic reflection
profiling: Journal of Geophysical Research, v. 85, pp. 4773-4800.
Bryan, K., 1909, Geology of the vicinity of Albuquerque: University of
New Mexico, Bulletin 51, Geological Series, v. 3, 24 pp.
Bryan, K., 1938, Geology and groundwater conditions of the Rio Grandedepression in Colorado and New Mexico; in U.S. Natural Resources
Planning Board, The Rio Grande Joint Investigations in the upper Rio
Grande Basin: U.S. Government Printing Office, Washington, D.C., v.
1, pt. 2, pp. 197-225.
Bryan, K., and McCann, F. T., 1937, The Ceja del Rio Puerco: A borderfeature of the Basin and Range province in New Mexico, Part I:
Journal of Geology, v. 45, pp. 801-828.
Bryan, K., and McCann, F. T., 1938, The Ceja del Rio Puerco: A borderfeature of the Basin and Range province in New Mexico, Part II:
Journal of Geology, v. 46, pp. 1-16.
Campbell, J. A., 1967, Geology and structure of a portion of the Rio
Puerco fault belt, western Bernalillo County, New Mexico:
Unpublished PhD dissertation, University of New Mexico, 89 pp.
Cather, S. M., 1989, Post-Laramide tectonic and volcanic transition inwest-central New Mexico: New Mexico Geological Society,
Guidebook 40, pp. 91-97.
Chamberlin, R. M., and Anderson, 0. J., 1989a, The Laramide Zuni uplift,
southeastern Colorado Plateau: a microcosm of Eurasian-styleindentation-extrusion tectonics: New Mexico Geological Society,
Guidebook 40, pp. 81-90.
Chamberlin, R. M., and Anderson, 0. J., 1989b, The southern margin of
the Colorado Plateau: a ragged edge bounded by Neogene domains ofcrustal extension: New Mexico Geological Society, Guidebook 40, pp.
37-38.
Chamberlin, R. M., Anderson, 0. J., Lucas, S. G., Maxwell, C. H., andLove, D. W., 1989, Second-day road log, from Grants to El Malpais,
Fence Lake, Zuni Pueblo and Gallup: New Mexico Geological Society,
Guidebook 40, pp. 25-48.
Clary, J. H., Korecki, N. T., and Mondragon, R. R., 1984, Geology ofAlbuquerque, New Mexico, United States of America: Bulletin of Association of Engineering Geologists, v. 21, no. 2, pp. 127-156.
Colpitts, R. M., Jr., Bauer, P., and Smith, C. T., 1989, Supplemental roadlog 4, from intersection of NM-53 and Forest Road 50 (near Bandera
Crater) to Grants, via Forest Roads 50 and 49: New Mexico GeologicalSociety, Guidebook 40, pp. 73-75.
Connolly, J. R., Woodward, L. A., and Hawley, J. W., 1982, Road-log
segment I-A: Albuquerque to Tijeras Canyon: New Mexico GeologicalSociety, Guidebook 33, pp. 2-8.
Cooley, M. E., and Akers, J. P., 1961, Ancient erosion cycles of the Little
Colorado River, Arizona and New Mexico: U.S. Geological Survey,
Professional Paper 424-C, pp. 244-248.
Cordell, Lindrith, 1978a, Gravity profile along Tramway road; in Hawley,
J. W. (compiler), Guidebook to Rio Grande rift in New Mexico andColorado: New Mexico Bureau of Mines & Mineral Resources,
Circular 163, 156 pp.
Cordell, Lindrith, 1978b, Regional geophysical setting of the Rio Grande
rift: Geological Society of America, Bulletin, v. 89, pp. 1073-1090.
Cordell, L. S., 1982a, Pottery Mound: New Mexico Geological Society,
Guidebook 33, pp. 42-43.
Cordell, L. S., 1982b, Prehistoric archaeology in Tijeras Canyon: New
Mexico Geological Society, Guidebook 33, pp. 8-9.
DeBrine, D. E., 1971, Quantitative hydrologic study of a closed basin witha playa, Estancia Valley, New Mexico: Unpublished PhD dissertation,
New Mexico Institute of Mining & Technology, 165 pp.

Dillinger, J. K., 1990, Geologic map of the Grants 30' X 60' quadrangle,
west-central New Mexico: U.S. Geological Survey, Coal Investigations
Map C-118-A, scale 1:100,000.
de Voogd, B., Brown, L. D., and Merey, C., 1986, Nature of the easternboundary of the Rio Grande rift from COCORP surveys in the Albuquerque basin, New Mexico: Journal of Geophysical Research, v. 91,
pp. 6305-6320.
Dutton, C. E., 1885, Mt. Taylor and the Zuni Plateau: U.S. Geological
Survey, 6th Annual Report.
Fenneman, N. M., 1931, Physiography of western United States: McGrawHill Book Co., New York, 534 pp.
Frye, J. C., Leonard, A. B., and Glass, H. D., 1982, Western extent of
Ogallala Formation in New Mexico: New Mexico Bureau of Mines &Mineral Resources, Circular 175, 41 pp.
Gallaher, B. M., and Goad, M. S., 1981, Water-quality aspects of uranium
mining and milling in New Mexico: New Mexico Geological Society,
Special Publication 10, pp. 85-91.
Galloway, R. W., 1970, The full-glacial climate in the southwestern
United States: Annals of the Association of American Geographers, v.
60, pp. 245-256.
Galusha, T., 1966, The Zia Sand Formation, new early to medial Miocene
beds in New Mexico: American Museum Novitates, no. 2271, 12 pp.
Galusha, T., and Buick, J. C., 1971, Stratigraphy of the Santa Fe Group,
New Mexico: American Museum of Natural History, Bulletin, v. 14, pp.
1-128.
Gardner, J. L., 1951, Vegetation of the creosote bush area of the Rio
Grande valley in New Mexico: Ecological Monographs, v. 21, pp. 379403.
Gawell, M. J., 1975, Chemical and petrographic variations in the Cerro
Negro-Cerrito Arizona undertone chain, Valencia [now Cibola] County,
New Mexico: Unpublished MS thesis, Kent State University, Kent,
Ohio, 57 pp.
Gawell, M. J., and Laughlin, A. W., 1975, Chemical and petrographic
variations in the Cerro Negro-Cerrito Arizona cinder cone Lineation,
Valencia [now Cibola] County, New Mexico (abs.): Geological Societyof America, Abstracts with Programs, v. 7, p. 1084.
Gawne, C. E., 1973, Faunas and sediments of the Zia Sand, middle
Miocene of New Mexico: Unpublished PhD dissertation, Columbia University, New York, 352 pp.
Gawne, C. E., 1981, Sedimentology and stratigraphy of the Miocene Zia
Sand of New Mexico: Geological Society of America, Bulletin, v. 92,
part I, pp. 999-1007, part II, p. 2484-2552.
Geissman, J. W., Brown, L., Turrin, B. D., McFadden, L. D., and Harlan,
S. S., 1990, Brunhes chron excursion/polarity episode recorded during
the late Pleistocene, Albuquerque Volcanoes, New Mexico, USA: International Geophysical Journal, v. 102, pp. 73-88.
Gile, L. H., Peterson, F. F., and Grossman, R. B., 1966, Morphologicaland genetic sequences of carbonate accumulation in desert soils: Soil
Science, v. 101, pp. 347-360.
Gile, L. H., Hawley, J. W., and Grossman, R. B., 1981, Soils and geomorphology in the Basin and Range area of southern New MexicoGuidebook to the Desert Project: New Mexico Bureau of Mines &Mineral Resources, Memoir 39, 222 pp.
Goddard, E. M., 1966, Geologic map and sections of the Zuni Mountains
fluorspar district, Valencia [now Cibola] County, New Mexico: U.S.
Geological Survey, Miscellaneous Investigations Series Map 1-454,
scale 1:31,680.
Goff, F., Gardner, J. N., Baldridge, W. S., Hulen, J. B., Nielson, D. L.,
Vaniman, D., Heiken, G., Dungan, M. A., and Braxton, D., 1989,
Excursion 17B: Volcanic and hydrothermal evolution of Pleistocene
Valles caldera and Jemez volcanic field; in Chapin, C. E., and Zidek, J.
(eds.), Field excursions to volcanic terranes in the western United States,
Volume I: Southern Rocky Mountain region: New Mexico Bureau of
Mines & Mineral Resources, Memoir 46, pp. 381-434.
Gordon, E. D., 1961, Geology and ground-water resources of the GrantsBluewater area, Valencia [now Cibola] County, New Mexico: New
Mexico State Engineer, Technical Report 20, 109 pp.
Green, M. W., and Jackson, T. J., 1976, Geologic and structure contour
maps of the Gallup East quadrangle, McKinley County, New Mexico:
U.S. Geological Survey, Open-File Report 76-453, scale 1:24,000.
Gustayson, T. C., and Winkler, D. A., 1988, Depositional facies of the
Miocene-Pliocene Ogallala Formation, northwestern Texas and eastern
New Mexico: Geology, v. 16, no. 3, pp. 203-206.
Grambling, J. A., and Wells, S. G. (eds.) 1982, Albuquerque country II:
New Mexico Geological Society, Guidebook 33, 370 pp.
Hack, J. T., 1942, Sedimentation and volcanism in the Hopi Buttes, Arizona: Geological Society of America, Bulletin, v. 53, pp. 335-372.

146
Hacker, L. W., 1977, Soil survey of Bemalillo County and parts of Sandoval and Valencia Counties, New Mexico: U.S. Soil Conservation
Service, 101 pp.
Hackman, R. J., and Olson, A. B., 1977, Geology, structure, and uranium
deposits of the Gallup 1° x 2° quadrangle, New Mexico and Arizona:
U.S. Geological Survey, Miscellaneous Investigations Series Map I981,
scale 1:250,000.
Harbour, J., 1958, Microstratigraphic and sedimentational studies of anearly man site near Lucy, New Mexico: Unpublished MS thesis, University of New Mexico, Albuquerque, 1 1 1 pp.
Harrington, E. R., 1934, The origin of ice caves: Journal of Geology, v.
43, pp. 433-436.
Hastings, J. R., and Turner, R. M., 1965, The changing mile: University of
Arizona Press, Tucson, 317 pp.
Hawley, J. W. (compiler) 1978, Guidebook to Rio Grande rift in New
Mexico and Colorado: New Mexico Bureau of Mines & Mineral Resources, Circular 163, 241 pp.
Hawley, J. W., 1984, The Ogallala Formation in eastern New Mexico; in
Whetstone, G. A. (ed.), Proceedings of the Ogallala Aquifer Symposium: Texas Tech University, Water Resources Center, pp. 157-175.
Hawley, J. W., 1986a, Physiographic provinces of New Mexico; in Williams, J. L. (ed.), New Mexico in maps, 2nd edition: University of New
Mexico, Albuquerque, pp. 23-27.
Hawley, J. W., 1986b, Environmental geology of the Keers
Environmental, Inc. Asbestos Disposal Site, Torrance County, New
Mexico: New Mexico Bureau of Mines & Mineral Resources, Open-file
Report 245, 12 pp., 5 figs.
Hawley, J. W., and Longmire, P. A. (in press), Site characterization andselection (Chapter 3); in Reith, C. C., and Thomson, B. M. (eds.),
Deserts as dumps: University of New Mexico Press, Albuquerque.
Hawley, J. W., Bachman, G. 0., and Manley, K., 1976, Quaternarystratigraphy in the Basin and Range and Great Plains provinces, New
Mexico and Texas; in Mahaney, W. C. (ed.), Quaternary stratigraphy ofNorth America: Dowden, Hutchinson, and Ross, Inc., Stroudsburg,
Pennsylvania, pp. 235-274.
Hawley, J. W., Crumpler, L. S., and Wells, S. G., 1982, Road-log segment
II-B: I-40/NM-6 interchange to LaJara Mesa northwest of Mount Taylor:
New Mexico Geological Society, Guidebook 33, pp. 49-70.
Hawley, J. W., Love, D. W., and Lambert, P. W., 1982, Road-log segment
I-C: Abo Canyon-Blue Springs area to Albuquerque via Belen and Los
Lunas: New Mexico Geological Society, Guidebook 33, pp. 24-30.
Hawley, J. W., Love, D. W., and Wells, S. G., 1982, Road-log segment
II-A: Albuquerque to Correo via El Cerro de Los Lunas and Rio Puerco:
New Mexico Geological Society, Guidebook 33, pp. 38-49.
Hawley, J. W., Love, D. W., and Wells, S. G., 1983, Summary of thehydrology, sedimentology, and stratigraphy of the Rio Puerco valley; in
Wells, S. G., Love, D. W., and Gardner, T. W. (eds.), Chaco Canyon
country: American Geomorphological Field Group, Field Trip Guidebook, pp. 33-36.
Hawley, J. W., Foster, R. W., Broadhead, R., and Love, D. W., 1982,
Road-log segment I-B: Tijeras Canyon to Abo Canyon via Estancia andManzano: New Mexico Geological Society, Guidebook 33, pp. 8-23.
Hawley, J. W., Lambert, P. W., Kelley, V. C., and Woodward, L. A.,
1982, Road-log segment III-B: Sandia Mountains at tramway terminal
to Placitas via Alameda, Rio Rancho, Ceja del Rio Puerco, Loma Duran,and Bemalillo: New Mexico Geological Society, Guidebook 33, pp. 8094.
Heath, D. L., 1983, Flood and recharge relationships of the lower Rio
Puerco: New Mexico Geological Society, Guidebook 34, pp. 329-337.
Heath, D. L., and Love, D. W., 1983, Multivariate historical processes
increasing thresholds for erosion within the Rio Puerco, north-central
New Mexico (abs.): Geological Society of America, Abstracts with
Programs, v. 18, no. 5, pp. 360.
Henderson, J., 1933, Caverns, ice caves, sinkholes, and natural bridges II:
Ice caves and related phenomena: University of Colorado Studies, v. 20,
pp. 115-158.
Herrick, C. L., 1904, Lake Otero, an ancient salt lake in southeastern New
Mexico: American Geologist, v. 34, pp. 174-189.
Hunt, C. B., 1974, Natural regions of the United States and Canada: W. H.
Freeman & Co., San Francisco, 725 pp.
Izett, G. A., Obradovich, J. D., Naeser, C. W., and Cebula, G. T., 1981,
Potassium-argon and fission-track zircon ages of Cerro Toledo Rhyolite
tephra in the Jemez Mountains, New Mexico: U.S. Geological Survey,
Professional Paper 1199-D, pp. 37-43.
Jenkins, D. N., 1982, Geohydrology of the Madera Group, western Estancia Basin, New Mexico: New Mexico Geological Society,
Guidebook 33, pp. 361-366.
Johnpeer, G., Robinson-Cook, S., Bobrow, D., Barrie, D., Kelliher, J.,

and McNeil, R., 1987, Geology and tunnelling, v. 3; in Lattman, L. H.,
and Roser, H. E. (authorized representatives), Estancia Basin, New
Mexico Superconducting Super Collider: New Mexico Bureau of Mines& Mineral Resources, Open-file Report 258, 224 pp.
Johnson, D. W., 1932, Rock fans of and regions: American Journal of
Science (5), v. 23, pp. 389-420.
Kaehler, C. A., 1990, Lithology of basin-fill deposits in the AlbuquerqueBelen basin, New Mexico: U.S. Geological Survey, Water-Resources
Investigations Report 89-4162, 14 pp.
Kasten, J. A., 1977, Petrology and geochemistry of calc-alkaline andesites
within the Albuquerque basin, Valencia County, New Mexico: Unpublished MS thesis, University of New Mexico, Albuquerque, 78 pp.
Kelley, V. C. (compiler) 1963, Geology and technology of the Grants
uranium region: New Mexico Bureau of Mines & Mineral Resources,
Memoir 15, 277 pp.
Kelley, V. C., 1972, Geology of the Fort Sumner sheet, New Mexico:
New Mexico Bureau of Mines & Mineral Resources, Bulletin 98, 55
pp.
Kelley, V. C., 1977, Geology of Albuquerque Basin, New Mexico: New
Mexico Bureau of Mines & Mineral Resources, Memoir 33, 59 pp.
Kelley, V. C., 1982, Albuquerque, its mountains, valley, water, and volcanoes: New Mexico Bureau of Mines & Mineral Resources, Scenic
Trips to the Geologic Past No. 9, 106 pp.
Kelley, V. C., and Kudo, A. M., 1978, Volcanoes and related basalts of the
Albuquerque basin, New Mexico: New Mexico Bureau of Mines &Mineral Resources, Circular 156, 30 pp.
Kelley, V. C., and Northrop, S. A., 1975, Geology of Sandia Mountains
and vicinity, New Mexico: New Mexico Bureau of Mines & Mineral
Resources, Memoir 29, 136 pp.
Kelley, V. C., and Wood, G. H., Jr., 1946, Lucero uplift, Valencia,
Socorro and Bemalillo Counties, New Mexico: U.S. Geological Survey,
Oil and Gas Investigations Preliminary Map 47.
Kelley, V. C., Woodward, L. A., Kudo, A. M., and Callender, J. F.,
1976, Guidebook to Albuquerque basin of the Rio Grande rift, New
Mexico: New Mexico Bureau of Mines & Mineral Resources, Circular
153, 30 pp.
Kelley, V. C., Hawley, J. W., and Wells, S. G., 1982, Road-log segment
III-A: Albuquerque to Sandia Peak Tramway via 1-25, Montgomery
Blvd., and Tramway Blvd.: New Mexico Geological Society,
Guidebook 33, pp. 76-80.
Kelly, T. E., and Reynolds, C. B., 1984, Structural geology of the Malpais
Valley, San Rafael, New Mexico: New Mexico Geological Society,
Guidebook 40, pp. 119-121.
Kemodle, J. W., and Scott, W. B., 1986, Three-dimensional model simulation of steady-state ground-water flow in the Albuquerque-Belenbasin, New Mexico: U.S. Geological Survey, Water-Resources Investigations Report 84-4353, 58 pp.
Kreitler, C. W., Mullican, W. F., and Nativ, R., 1990, Hydrogeology of the
Diablo Plateau, Trans-Pecos Texas; in Hydrogeology of Trans-Pecos
Texas: The University of Texas at Austin, Bureau of Economic
Geology, Guidebook 25, pp. 49-58.
Kudo, A. M., Kelley, V. C., Damon, P. E., and Shafiqullah, M., 1977, KAr ages of basalt flows at Canjilon Hill, Isleta volcano, and the Cat Hills
volcanic field, Albuquerque-Belen, central New Mexico: Isochron/West, no. 18, pp. 15-16.
Lagasse, P. F., 1981, Geomorphic response of the Rio Grande to dam
construction; in Wells, S. G., and Lambert, W. (eds.), Environmental
geology and hydrology in New Mexico: New Mexico Geological Society, Special Publication 10, pp. 27-46.
Lambert, P. W., 1968, Quaternary stratigraphy of the Albuquerque area,
New Mexico: Unpublished PhD dissertation, University of New
Mexico, 329 pp.
Lambert, P. W., Hawley, J. W., and Wells, S. G., 1982, Supplemental
road-log segment III-S: Urban and environmental geology of the Albuquerque area: New Mexico Geological Society, Guidebook 33, pp.
97-119.
Laughlin, A. W., Aldrich, M. J., Jr., Ander, M. E., Heiken, G. H., andVaniman, D. T., 1982, Tectonic setting and history of late Cenozoic
volcanism in west-central New Mexico: New Mexico Geological Society, Guidebook 33, pp. 279-284.
Leavy, B. D., and Shafiqullah, M., 1987, New K-Ar analyses of basalts
from southern California and central New Mexico: Isochron/West, no.
48, pp. 19-20.
Leavy, D. G., Phillips, F. M., Elmore, D., Kubick, P. W., and Gladney, E.,
1987, Measurement of cosmogenic 36Cl/Cl in young volcanic rocks: anapplication of accelerator mass spectrometry in geochronology: Nuclear
Instruments and Methods in Physics Research, B26, pp. 246-250.

147
Lee, W. T., 1926, An ice cave in New Mexico: Geographical Review, v.
16, pp. 55-59.
Leopold, L. B., 1946, Two intense local floods in New Mexico: American
Geophysical Union, Transactions, v. 27, pp. 535-539.
Leopold, L. B., 1951, Pleistocene climate in New Mexico: American
Journal of Science, v. 249, pp. 152-168.
Leopold, L. B., and Snyder, C., 1951, Alluvial fills near Gallup, New
Mexico: U.S. Geological Survey, Water-Supply Paper 1110-A, 19 pp.
Lindsey, A. A., 1951, Vegetation and habits in a southwestern volcanicarea: Ecological Monographs, v. 21, pp. 227-253.
Lipman, P. W., and Mehnert, H. H., 1980, Potassium-argon ages from theMount Taylor volcanic field, New Mexico: U.S. Geological Survey,
Professional Paper 1124-B, 8 pp.
Love, D. W., 1986, A geological perspective of sediment storage anddelivery along the Rio Puerco, central New Mexico; in Hadley, R. F.
(ed.), Drainage basin sediment delivery: International Association of
Hydrological Sciences, Publication 159, pp. 305-322.
Love, D. W., 1989a, Geomorphic development of the Rio San Jose valley:
New Mexico Geological Society, Guidebook 40, pp. 11-12.
Love, D. W., 1989b, Bidahochi Formation: An interpretive summary: New
Mexico Geological Society, Guidebook 40, pp. 273-280.
Love, D. W., and Gillam, M. (in press), Navajo and Acoma-Zuni sections;in Morrison, R. B. (ed.), Quaternary non-glacial geology; conterminous
U.S.: Geology of North America, v. K-2, Geological Society ofAmerica, Boulder, Colorado.
Love, D. W., and Hadley, R. F., 1986, Second day, road log from Albuquerque to Rio Puerco, Correo, Los Lunas, Belen, and Bernardo, andretum to Albuquerque; in Hadley, R. F. (ed.), International symposium
on drainage basin sediment delivery, program of paper sessions and roadlogs of field trips: University of Denver, Denver, Colorado, pp. 1-9.
Love, D. W., and Young J. D., 1983, Progress report on the lateCenozoic evolution of the lower Rio Puerco: New Mexico GeologicalSociety, Guidebook 34, pp. 277-284.
Love, D. W., Hawley, J. W., and Young J. D., 1982, Preliminary report on
the geomorphic history of the lower Rio Puerco in relation to archeological sites and cultural resources of the lower Hidden Mountain
Dam Site; in Eidenbach, P. (ed.), Inventory survey of the lower Hidden
Mountain floodpool, lower Rio Puerco drainage, central New Mexico:
Human Systems Research, Tularosa, New Mexico, pp. 20-65. Lozinsky,
R. P., 1988, Stratigraphy, sedimentology, and sand petrology of the
Santa Fe Group and pre-Santa Fe Tertiary deposits in the Albuquerquebasin, central New Mexico: Unpublished PhD dissertation, New Mexico
Institute of Mining & Technology, 298 pp.
Lozinsky, R. P., and Hawley, J. W., 1986, Upper Cenozoic Palomas
Formation of south-central New Mexico: New Mexico Geological Society, Guidebook 37, pp. 239-248.
Lucas, S. G., and Anderson, 0. J., 1989, Supplemental road log 2, fromBlack Rock to Ojo Caliente: New Mexico Geological Society, Guidebook 40, pp. 69-70.
Lyons, T. R., 1969, A study of the paleo-Indian and desert culture complexes of the Estancia Valley area, New Mexico: Unpublished PhD
dissertation, University of New Mexico, Albuquerque, 355 pp.
Machette, M. N., 1978a, Geologic map of the San Acacia quadrangle,
Socorro County, New Mexico: U.S. Geological Survey, Geologic Quadrangle Map GQ-14l5, scale 1:24,000.
Machette, M. N., 1978b, Preliminary geologic map of the Socorro 1° by 2°
quadrangle, central New Mexico: U.S. Geological Survey, Open-file
Report 78-607.
Machette, M. N., 1978c, Dating of Quaternary faults in the southwestern
United States by using buried calcic paleosols: U.S. Geological Survey,
Journal of Research, v. 6, pp. 369-381.
Machette, M. N., 1982, Quaternary and Pliocene faults in the La Jencia
and southern part of the Albuquerque-Belen basins, New Mexico: Evidence of fault history from fault-scarp morphology and Quatemary geology: New Mexico Geological Society, Guidebook 33, pp. 161-169.
Machette, M. N., 1985, Calcic soils of the southwestem United States:
Geological Society of America, Special Paper 203, pp. 1-21.
Machette, M. N., and McGimsey, R. G., 1983, Map of Quaternary and
Pliocene faults in the Socorro and western part of the Fort Sumner 1° x
2° quadrangles, central New Mexico: U.S. Geological Survey, Miscellaneous Field Studies Map MF-1465-A, scale 1:250,000, 12 pp. pamphlet.
Maxwell, C. H., 1982, El Malpais: New Mexico Geological Society,
Guidebook 33, pp. 299-301.
Maxwell, C. H., 1986, Geologic map of El Malpais lava field and surrounding areas, Cibola County, New Mexico: U.S. Geological Survey,
Miscellaneous Investigations Map 1-1595, scale 1:62,500.
Maxwell, C. H., Anderson, 0. J., Lucas, S. G., Chamberlin, R. M., andLove, D. W., 1989, First-day road log, from Albuquerque to Mesita,

Laguna, Acoma, McCartys and Grants: New Mexico GeologicalSociety, Guidebook 40, pp. 1-24.
McCann, F. T., 1938, Ancient erosion surface in the Gallup-Zuni area,
New Mexico: American Journal of Science (5), v. 36, pp. 260-279.
McGrath, D. B., and Hawley, J. W., 1987, Geomorphic evolution andsoil-geomorphic relationships in the Socorro area, central New Mexico;
in McLemore, V. T., and Bowie, M. R. (compilers), Guidebook to theSocorro area, New Mexico: New Mexico Bureau of Mines & Mineral
Resources, pp. 55-67.
McLellan, M. W., Robinson, L. N., Haschke, L. R., Carter, M. D., andMedlin, A., 1982, Fence Lake Formation (Tertiary), west-central New
Mexico: New Mexico Geology, v. 4, pp. 53-55.
McQuillan, D. M., 1982, Pollution of the Rio Grande valley-fill aquifer:
New Mexico Geological Society, Guidebook 33, pp. 357-360. Meinzer,
0. E., 1911, Geology and water resources of Estancia Valley, New
Mexico: U.S. Geological Survey, Water-Supply Paper 275, 89 pp.
Meinzer, 0. E., 1922, Map of the Pleistocene lakes of the Basin andRange province and its significance: Geological Society of America,
Bulletin, v. 33, pp. 541-552.
Meinzer, 0. E., and Hare, R. E., 1915, Geology and water resources of theTularosa Basin, New Mexico: U.S. Geological Survey, Water-Supply
Paper 343, 317 pp.
Miller, J. R., and Wells, S. G., 1986, Types and processes of short-term
sediment and uranium-tailings storage in arroyos: an example from the
Rio Puerco of the West, New Mexico; in Hadley, R. F. (ed.), Drainage
basin sediment delivery: International Association of Hydrological Sciences, Publication 159, pp. 335-353.
Moench, R. H., and Schlee, J. S., 1967, Geology and uranium deposits of
the Laguna district, New Mexico: U.S. Geological Survey, Professional
Paper 519, 117 pp.
Myers, D. A., 1977, Geologic map of the Scholle quadrangle, Socorro,
Valencia, and Torrance counties, New Mexico: U.S. Geological Survey,
Geologic Quadrangle Map GQ-1412, scale 1:24,000.
Myers, D. A., and McKay, E. J., 1976, Geologic map of the north end of
the Manzano Mountains, Tijeras and Sedillo quadrangles, Bemalillo
County, New Mexico: U.S. Geological Survey, Miscellaneous Geologic
Investigations Map 1-968.
Myers, D. A., McKay, E. J., and Sharps, J. A., 1981, Geologic map of
the Becker quadrangle, Valencia and Socorro counties, New Mexico:
U.S. Geological Survey, Geologic Quadrangle Map GQ-1556.
Nichols, R. L., 1946, McCartys basalt flow, Valencia [Cibola]
County, New Mexico: Geological Society of America, Bulletin, v.
57, pp. 10491066.
O'Neill, F. M., and Rigby, J. K., Jr., 1982, A rare fossil skeleton of Camelops from Pleistocene deposits near Albuquerque: New Mexico
Geological Society, Guidebook 33, pp. 82-84.
Phillips, F. M., Leavy, B. D., Jannik, N. 0., Elmore, D., and Kubik, P. W.,
1986, The accumulation of cosmogenic chlorine-36 in rocks: a methodof surface exposure dating: Science, v. 231, pp. 41-43.
Popp, C. J., Love, D. W., Hawley, J. W., and Novo-Gradac, K., 1984,
Radionuclide and heavy metal distribution in 20th century sediments ofmajor streams in the eastern part of the Grants uranium region, New
Mexico: New Mexico Bureau of Mines & Mineral Resources, Hydrologic Report 7, pp. 34-48.
Popp, C. J., Hawley, J. W., Love, D. W., and Dehn, M., 1988, Use ofradiometric (Cs-137, Pb-210), geomorphic, and stratigraphic techniques
to date recent oxbow sediments in the Rio Puerco drainage, Grants
uranium region, New Mexico: Environmental Geology and Water Science, v. 11, no. 3, pp. 253-269.
Pearce, T. M., 1965, New Mexico place names, a geographical dictionary:
University of New Mexico, Albuquerque, 187 pp.
Rautman, C. A. (compiler) 1980, Geology and mineral technology of theGrants uranium region 1979: New Mexico Bureau of Mines & Mineral
Resources, Memoir 38, 400 pp.
Reiche, P., 1949, Geology of the Manzanita and north Manzano Mountains, New Mexico: Geological Society of America, Bulletin, v. 60, pp.
1183-1212.
Robertson, J. F., 1975, Preliminary geologic map of the Pinedal,
McKinley County, New Mexico: U.S. Geological Survey, Open-File
Report 74244, scale 1:24,000.
Robinson, S., and Love, D. W., 1986, Applications of detailed sedimentologic studies of modern waning-flow deposits to alluvial stratigraphy(abs.): Geological Society of America, Abstracts with Programs, v. 18,
no. 5, p. 407.
Schumm, S. A., 1977, The fluvial system: John Wiley & Sons, New York,238 pp.
Shepherd, R. G., 1976, Sedimentary processes and structures of ephemeral
stream point bars, Rio Puerco west of Albuquerque, New Mexico (abs.):

148
Geological Society of America, Abstracts with Programs, v.
8, no. 6, p. 1103.
Shepherd, R. G., 1987, Lateral accretion surfaces in ephemeralstream point bars, Rio Puerco, New Mexico; in Ethridge, F. G.,
Flores, R. M., and Harvey, D. D. (eds.), Recent developments influvial sedimentology: Society of Economic Paleontologists andMineralogists, Special Publication 39, pp. 93-98.
Slack, P. B., and Campbell, J. A., 1976, Structural geology of the Rio
Puerco fault zone and its relationship to central New Mexico tectonics:
New Mexico Geological Society, Special Publication 6, pp. 46-52.
Smith, R. E., 1957, Geology and ground-water resources of Torrance
County, New Mexico: New Mexico Bureau of Mines & Mineral Resources, Ground-Water Report 5, 186 pp.
Smith, R. L., Bailey, R. A., and Ross, C. S., 1970, Geologic map of theJemez Mountains, New Mexico: U.S. Geological Survey, Miscellaneous
Investigations Map I-571.
Smith, L. N., and Anderson, R. Y., 1982, Pleistocene-Holocene climateof the Estancia Basin, central New Mexico: New Mexico GeologicalSociety, Guidebook 33, pp. 347-350.
Smith, L. N., Bullard, T. F., and Wells, S. G., 1982, Quaternary geology
and geomorphology of Tijeras Canyon, New Mexico: New Mexico
Geological Society, Guidebook 33, pp. 5-7.
Spiegel, Z., and Baldwin, B., 1963, Geology and water resources of the
Santa Fe area, New Mexico: U.S. Geological Survey, Water-Supply
Paper 1525, 158 pp.
Stark, J. T., and Dapples, E. C., 1946, Geology of the Los Pinos Mountains, New Mexico: Geological Society of America, Bulletin, v. 57, pp.
1121-1172.
Stark, J. T., 1956, Geology of the south Manzano Mountains, New
Mexico: New Mexico Bureau of Mines & Mineral Resources, Bulletin
34, 49 pp.
Stone, W. J., 1984, Localization of fresh ground-water bodies-a special
consideration in siting landfills along the Rio Grande valley: New
Mexico Bureau of Mines & Mineral Resources, Hydrologic Report 7,
pp. 227238.
Stearns, C. E., 1953, Tertiary geology of the Galisteo-Tonque area, New
Mexico: Geological Society of America, Bulletin, v. 64, pp. 459-508.
Tedford, R. H., 1981, Mammalian biochronology of the later Cenozoic
basins of New Mexico: Geological Society of America, Bulletin, v. 92,
part I, pp. 1008-1022.
Tedford, R. H., 1982, Neogene stratigraphy of the northwestern Albuquerque basin: New Mexico Geological Society, Guidebook 33, pp.
273-278.
Thomson, B. M., and McQuillan, D. M., 1984, Nitrate contamination ofground water in Albuquerque: New Mexico Bureau of Mines &Mineral Resources, Hydrologic Report 7, pp. 204-216.
Titus, F. B., Jr., 1961, Ground-water geology of the Rio Grande troughin north-central New Mexico, with sections on the Jemez caldera
and the Lucero uplift: New Mexico Geological Society, Guidebook
12, pp. 186-192.
Titus, F. B., Jr., 1963, Geology and ground-water conditions in eastern
Valencia County, New Mexico: New Mexico Bureau of Mines &Mineral Resources, Ground-Water Report 7, 113 pp.
Titus, F. B., Jr., 1969, Late Tertiary and Quaternary hydrogeology of
Estancia basin, central New Mexico: Unpublished PhD dissertation,

University of New Mexico, Albuquerque, 179 pp.
Titus, F. B., Jr., 1973, Hydrogeologic evolution of Estancia Valley, aclosed basin in central New Mexico: New Mexico Bureau of Mines &Mineral Resources, Open-File Report 69, 184 pp.
Titus, F. B., Jr., 1980, Ground water in the Sandia and northern
Manzano Mountains, New Mexico: New Mexico Bureau of Mines &Mineral Resources, Hydrologic Report 6, 66 pp.
U.S. Army Corps of Engineers (compilers) 1979, Albuquerque Greater
Urban Area Studies Program: U.S. Government Printing Office,Washington, D.C., 13 volumes.
Van Devender, T. R., 1990, Late Quaternary vegetation and climate of
the Chihuahuan Desert, United States and Mexico; in Betancourt, J.
L., Van Devender, T. R., and Martin, P. S. (eds.), Fossil packrat
middens: The last 40,000 years of biotic change: University of
Arizona Press, Tucson, pp. 104-133.
Vazzana, M. E., and Ingersoll, R. V., 1981, Stratigraphy,
sedimentology, petrology, and basin evolution of the Abiquiu
Formation (Oligo-Miocene), north-central New Mexico: Summary:Geological Society of America, Bulletin, v. 91, Part I, pp. 990-992.
Weide, D. L., Schneider, G. B., Mytton, J. W., and Scott, G. R., 1979,
Geologic map of the Pueblo Bonito quadrangle, San Juan County, New
Mexico: U.S. Geological Survey, Miscellaneous Field Studies Map MF1119, scale 1:24,000.
Wells, S., Bullard, T. F., Condit, C. D., Jercinovic, M., Lozinsky, R. P.,
and Rose, D. E., 1982, Geomorphic processes on the alluvial valley
floor of the Rio Puerco: New Mexico Geological Society, Guidebook
33, pp. 45-47.
Wells, S. G., Love, D. W., and Gardner, T. W. (eds.) 1983, Chaco
Canyon country, a field guide to the geomorphology, Quaternary
geology, paleoecology, and environmental geology of northwestern
New Mexico: American Geomorphological Field Group, Field Trip
Guidebook, 253 pp.
White, W. D., 1989, Geohydrologic and environmental indicators of adewatered wetland: Ojo del Gallo, San Rafael, New Mexico: New
Mexico Geological Society, Guidebook 40, pp. 337-345.
Wilson, L., 1981, Potential for ground-water pollution in New Mexico;
in Wells, S. G., and Lambert, W. (eds.), Environmental geology andhydrology in New Mexico: New Mexico Geological Society, Special
Publication 10, pp. 47-54.
Wood, W. W., and Sanford, W. E., 1990, Ground-water control ofevaporite deposition: Economic Geology, v. 85, pp. 1226-1235.
Woodward, L. A., Callender, J. F., Seager, W. R., Chapin, C. E., Gries,
J. C., Shaffer, W. L., and Zilinsky, R. E., 1978, Tectonic map of the
Rio Grande rift region in New Mexico, Chihuahua, and Texas, Sheet
2; in Hawley, J. W. (compiler), Guidebook to Rio Grande rift in New
Mexico and Colorado: New Mexico Bureau of Mines & Mineral Resources, Circular 163.
Wright, H. E., Jr., 1946, Tertiary and Quaternary geology of the lower
Rio Puerco area, New Mexico: Geological Society of America, Bulletin,v. 57, pp. 383-456.
Young, J. D., 1982, Late Cenozoic geology of the lower Rio Puerco,
Valencia and Socorro Counties, New Mexico: Unpublished MS thesis,
New Mexico Institute of Mining & Technology, Socorro, New Mexico,
126 pp.

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Laminated ice bodies in collapsed lava tubes at El Malpais National Monument, central New Mexico
Lonnie Thompson', Ellen Mosley-Thompson', Julio L. Betancourt², David W. Love', Alex Wilson', Greg Lecnard4,
and R. Scott Anderson''Byrd Polar Research Institute, Ohio State University, Columbus, Ohio 43210-1308; ²U.S. Geological Survey, 1675 W. Anklam Rd., Tucson, Arizona 85705;'New Mexico Bureau of Mines & Mineral Resources, Socorro, New Mexico 87801; °Department of Geosciences, University of Arizona, Tucson, Arizona 85721;'Quaternary Studies Program, Northern Arizona University, Flagstaff, Arizona 86011

Ice deposits in caves have a global distribution, primarily
in high-elevation or high-latitude lava tubes (Henderson,
1933). Unlike ice sheets, cave ice has been virtually ignored
as a climate proxy. Laminated ice bodies from collapsed
lava tubes at El Malpais National Monument, central New
Mexico (Fig. 1; Lee, 1926; Harrington, 1934; Lindsey,
1951; Elston and Wohletz, 1987) are now being evaluatedfor paleoclimatic value. The ice-cave systems can bethought of as frozen and very local, perched water tables. In
June 1990, three cores were extracted from Candelaria and
La Marchanita Caves. These cores are now being dated andanalyzed for pollen, microparticle concentrations, oxygenisotope ratios, and NO-3, Cl-, SO4-2, and PO4.

Our preliminary results suggest that the ice laminations (Fig.
1) are not annual, but instead represent occasional layers
between periods of melt. While interpretation of these data is
still evolving, measurements made at all depths within the
cores vary greatly between laminations. Opportunities exist to
calibrate the older parts of the ice body against ice
accumulated entirely within the 20th century. This younger ice
has refilled an area excavated by U.S. Army Cavalry stationed
at nearby Ft. Wingate beginning ca. 1850. Radiocarbon datesof the older ice deposits were obtained using accelerator mass
spectrometry. At Candelaria Cave, we dated a twig enclosed
near the bottom of the 4 m bank of 1 aminated ice to 3166 ±77 yrs B.P. Direct dates on the CO 2 gas within the ice ranged
from 1857 ± 56 yrs B.P. near the bottom to near modern at the
top. More work in the more accessible deposits, such as
Candelaria and La Marchanita Caves, could provide basic
approaches to develop the paleoclimatic worth of these
deposits and of cave ice in other parts of the world.
References
Elston, W. E., and Wohletz, K. H., 1987, Quaternary basalt fields of westcentral New Mexico: McCartys pahoehoe flow, Zuni Canyon as flow,
Auni Ice Cave, Bandera Crater, and Zuni Salt Lake Maar: Rocky
Mountain Section of the Geological Society of America, Cent( nnial
Field Guide, v. 2, pp. 431-436.
Harrington, E. R., 1934, The origin of ice caves: Journal of neology, v. 43,
pp. 433-436.
Henderson, J., 1933, Caverns, ice caves, sinkholes, and natural bridges II:
Ice caves and related phenomena: The University of Colorado Studies,v. 20, pp. 115-158.
Lee, W. T., 1926, An ice cave in New Mexico: Geographical Review, v.
16, pp. 55-59.
Lindsey, A. A., 1951, Vegetation and habitats in a southwestern volcanicarea: Ecological Monographs, v. 21, pp. 227-253.

Late Cenozoic geomorphic and tectonic evolution of the
Rio San Jose and tributary drainages within theBasin and Range/Colorado Plateau transition zone in west-central New Mexico
P. G. Drake', C. D. Harrington', S. G. Wells', F. V. Perry', and A. W. Laughlin'¹Glorieta Geoscience Inc., P.O. Box 5727, Santa Fe, New Mexico 87502;
Earth and Space Sciences Division, Los Alamos National Laboratory, Los Alamos, New Mexico 87545;'Department of Geology, University of New Mexico, Albuquerque, New Mexico 87131

2

Introduction
PurposeThis paper examines the late Cenozoic stratigraphy andgeologic history of the Rio San Jose and tributary drainages
between Grants and Laguna, New Mexico. The focus of
this research is to document the timing and spatial variationin tectonism/uplift across the Colorado Plateau/Basin andRange transition zone in west-central New Mexico during
the late Cenozoic.
Setting
The Rio San Jose originates at the Continental Divide on
the Colorado Plateau and flows eastward for a distance ofapproximately 185 km before emptying into the Rio Puerco
within the Rio Grande rift. The Rio San Jose drain s an areaof 9700 km' (Wright, 1946; Love, 1989), including parts of
the Zuni Mountains, Mt. Taylor, the Zuni—Bandera volcanic field, and the Lucero uplift. The portion o f the Rio
San Jose examined in this study lies within the :ransitionzone between the tectonically active Rio Grande rift and theless tectonically active Colorado Plateau (Fig. 1). The study
area extends from the Zuni uplift across the Acoina sag tothe edge of the Lucero uplift—Puerco fault zone (Fig. 2)
and lies athwart the Jemez lineament.
The bedrock stratigraphy within the study area consistsof nearly flat-lying Mesozoic sedimentary rocks overlain
unconformably by Pliocene and Quaternary volcanic rocks

which form high mesas located 200 to 300 m above local
base level. These high mesa basalts and associated volcanic
rocks of the Mt. Taylor area were erupted during a geologically brief period between 4 and 1.5 m.y. (Lipman andMehnert, 1979; Crumpler, 1982; Perry et al., 1990). Younger
Pleistocene and Holocene basalt flows lie at the level of themodern valley floor along the Rio San Jose (Lipman andMehnert, 1979; Maxwell, 1982; Perry and Laughlin, unpubl .
data).
Previous work
Early work on the geology of the Rio San Jose drainage
basin, by Bryan and McCann (1938) and Wright (1946),
focused on the geomorphic history of the Llano de Albuquerque and the lower Rio Puerco basin. Bryan and McCann(1938) suggested that the erosion surface underlying Mt.
Taylor was correlative with the late Tertiary regionally extensive Ortiz surface of the central rift. This hypothesis was
supported by Bachman and Mehnert (1978) through the useof K—Ar dates on basalt flows in the rift and in the Mt.
Taylor area. Wright (1946) defined the Mush Mesa and
Suwanee surfaces which are inset below the Ortiz surface
along the lower Rio San Jose.
Several recent studies have considered aspects of the geomorphic history of the Rio San Jose and tributary drainages.
Moench and Schlee (1967) defined the Wheat Mountainsurface, a Pliocene pediment surface overlain by basalt flows
on Frog and Clay Mesas northwest of Laguna (Fig. 1) whichis inset below, and is much less extensive than, the Ortiz
surface. Love (1989) suggests that the Rio San Jose drainage
basin developed in late Oligocene time and has evolved in
response to subsidence of the Albuquerque basin to the eastand episodic uplift of the Colorado Plateau to the west.

151
Maximum valley incision occurred between 1.5 and 0.38
m.y. (Love, 1989). Geomorphic surfaces in Lobo Canyon, atributary to the Rio San Jose which drains westward from
Mt. Taylor, have been described by Hunt (1938) and
Grimm (1985).
Pliocene and Quaternary geology
Stratigraphy
The Plio-Quaternary stratigraphy in the Grants/Laguna
area consists of two Pliocene and five Quaternary units thatlie above the modern flood plain or arroyo floor (Fig. 3).
Pliocene deposits (Tpal 1 and Tpal2) underlie erosion surfaces Tpl and Tp2 and are limited to pediment and axial
gravel deposits (Table l) which cap geomorphic surfaces
cut on Cretaceous sedimentary rocks. Tpal 1 and Tpal2 are
generally coarse-grained and contain greater proportions of
granite and quartzite clasts than are observed in Quaternary
deposits (Table l). Tpal 1 includes both a granite- andquartzite-dominated axial facies and a volcanic-dominatedtributary facies. Associated fine-grained deposits are
reddened but do not show evidence of pedogenesis.
Quaternary sediments (Qal1—Qal5) are dominated by fan
and terrace deposits, with pediment gravel limited to isolated
Qal1 deposits (Table l). Quaternary deposits contain a predominance of volcanic clasts, primarily basalt, porphyritic
andesite, and latite or quartz latite. Qal4 and Qal5 deposits are
generally comprised of fine-grained sand, silt, and clay which
form the modern valley fill, whereas preserved remnants ofolder Quaternary deposits (Qal3, Qal2, and Qal1) are coarser,
comprised primarily of coarse-grained gravelly sediments.
Drill logs for wells in the Laguna and Grants areas indicatethat 40 to 60 m of interbedded basalt and alluvium are present
below the level of the modern valley floor along the Rio San
Jose drainage (Gordon, 1961; Risser and Lyford, 1983),
contrasting with a 20 m maximum observed (preserved)
thickness of Qal1—Qal3.

Soils
Carbonate soils have developed in the semiarid climateof the Grants—Laguna area, and soil carbonate
morphology and Bt horizon development are themorphologic characteristics which best distinguish thesesoils. Soils in the study area exhibit stage I through III
carbonate morphology (Table 2) (terminology after Gile etal., 1981, and Machete, 1985). Bt horizons exhibit amaximum 5YR color, strong angular blocky or prismatic
structure, and thick, continuous clay films (Table 2). Soils
formed in Pliocene deposits and Qal1 are commonly
stripped, lacking carbonate and B1 horizons (Table 2).

Spatial variation of surficial deposits andgeomorphic surfaces
The high Pliocene pediment surface (Tp l) occ Hs as extensive mesas throughout the study area, whereas s Quaternary surfaces older than the valley floor are preserved only
as isolated remnants, primarily in tributary drainage basins(Fig. 4). Pliocene surfaces generally slope away from Mt.
Taylor or toward the Rio San Jose with gradients of 1-4°.
Quaternary surfaces located above the valley f oor (Q1—
Q3) are highly dissected. The Q1 surfaces include both anextensive pediment surface located along the upper Arroyo
del Valle drainage and a large fan-surface remnant preservedat the mouth of Water Canyon (Fig. 4). Q2 surfaces arepreserved primarily as surfaces cut on terrace deposits along
tributary drainages and along the Rio San Jose near Old
Laguna (Fig. 5). Q3 includes surfaces underlain try fan andterrace deposits located along tributary streams, and surfaces
associated with erosional terrace deposits along the Rio San
Jose (Fig. 4). Q4 surfaces consist of locally preserved proximal fans and terraces at, or less than, 10 m above themodern valley floor. Q5 terrace and fan surface< form thevalley floor along the Rio San Jose and tributary drainages.

Terrace heights and longitudinal profiles
The height of surfaces along the Rio San Jose and tributary streams ranges from 5 to 300 m above local baselevel. Pliocene surfaces lie between 200-300 m (Tp 1) and
140180 m (Tp2) above local base level. Quaternary surfaces
are located within 90 m, and Holocene surfaces within 5 m,of local base level (Table 3). Notably, relative elevations ofcorrelative surfaces increase both to the north and east
(Table 3). A distinct nickpoint in the Rio San Jose longprofile occurs between Encinal Creek and Old Laguna,where the river has downcut through the Laguna flow (asnamed by Lipman) (Fig. 5).

Pliocene and Quaternary faults
Several Pliocene and two Quaternary faults are recognizedwithin the study area. The Pliocene faults exhibit apredominant north to northeast trend throughout the Mount
Taylor volcanic field (Perry et al., in press). Structures offsetting Pliocene basalts were mapped on Grants Ridge, East
Grants Ridge, and Horace Mesa by Thaden et al. (1967) andbetween Rinconada and San Jose Canyons by Lipman et al.
(1979). A group of faults with offsets ranging from 14 to 60
m defines a series of northeast-trending horsts and grabens
along the southern end of Horace Mesa (Thaden et al. ,1967; Maxwell, 1977; Drake, unpubl. data). In contrast, Tp2
surfaces are offset 1.5 and 3 m on Black Mesa and Grants
Ridge (Thaden et al., 1967; Drake, unpubl. data). Faults do
not offset the volcanic rocks that comprise the Mt. Taylor
volcanic cone (Perry et al., 1990).
Possible Quaternary faults were identified at Cañon Seama
and one mile north of Old Laguna. A small graben appears
to offset the buried Laguna flows by 4 to 6 m at Callon
Seama just east of Flower Mountain (Fig. 5), based on drill
logs from Risser and Lyford (1983). A subvertical fault
which offsets Quaternary fluvial sediments a minimum of 6
m is exposed in a roadcut on Highway 279 approximately
one mile north of Old Laguna.

Results
Methods for estimating deposit and surface ages
The estimated ages of surficial deposits in the Grants/
Laguna area are based on stratigraphic and geomorphic position with respect to K—Ar-dated basalt flows and on comparison to radiometrically dated soil profiles developed inalluvium under similar climatic conditions. Using the statefactor approach of Jenny (1941), ages of soil profiles may
be estimated by comparison with similarly developed soil
profiles for which absolute-age determinations areavailable, if climate, topography, vegetation, and parent
material are similar (Harden, 1982; Birkeland, 1984). Soils
used for comparison in this study are located in the
Albuquerque basin, on the Llano de Albuquerque, and in the
Española basin.
Stratigraphic and geomorphic evolution of
the Grants—Laguna area
The Pliocene and Quaternary deposits and associated geomorphic surfaces in the Grants—Laguna area record a historyof long-term incision interrupted by relatively brief periods ofbase-level stability and/or aggradation. In contrast to theoverall Plio-Quaternary history, the late Quaternary has beencharacterized by alternating periods of erosion and deposition. The 200-300 m of incision into Tp1 between 2.5 m.y.
and the present time (using surface age estimated in Table 4)
is probably a response to regional uplift, whereas

periods of erosion and deposition during the late Quaternary
may be in response to climatic fluctuations, episodic volcanism, or differential uplift.
Tpal 1 was deposited as a veneer capping an extensive
pediment surface formed between 4 and 2.5 m.y. (Table 4)during a period of regional base-level stability and pedimentation in west-central New Mexico. Tp1 surfaces, which
are correlative with the Ortiz surface of the rift, were graded toan ancestral Rio San Jose that flowed near its present course.
The eruption of Mt. Taylor volcanic rocks onto Tp1 began by
3.73 ± 0.09 m. y. (Perry et al., 1990) and continued until
incision of Tp1 began at ca. 2.5 m.y.
Rapid incision of valleys followed abandonment of Tp
1, with approximately 60 m of incision between ca. 2.5
and 2.4 ± 0.18 m.y. (based on ages of flows capping Tp1
and Tp2 surfaces). During Tpal2 deposition the Rio San
Jose was a high-energy stream transporting coarse gravelfrom the Zuni Mountains and the Mt. Taylor area. Fluvial
processes during Tpal2 deposition were dominated bylateral planation and pedimentation.
Regional incision resumed along the Rio San Jose at ca.
2.4 m.y. and persisted into early—middle Pleistocene time(from 2.4 to 0.7-0.8 m.y.). This was a time of relatively
continuous base-level fall, evidenced by the absence ofearly Pleistocene surfaces, with 50-100 m of incision
between the time of Tp2 and Q 1. The formation of the Q1
surface marks the end of regional base-level fall during theearly Pleistocene.

Deep incision into the Mt. Taylor volcanic cone by Water
Creek began during Qal1 deposition, producing a large fan
at the mouth of Water Canyon. During Q1 time stream
systems located east of Water Creek were forming strath
terrace and pediment surfaces. The concurrent formation ofdepositional and erosional surfaces suggests that: (1) initialbreaching of the Mount Taylor amphitheater provided an
excess sediment load to Water Creek and resulted in deposition of the Water Canyon fan, or (2) this may have beena period of differential uplift with the Water Canyon fan
forming in a structural depression (the Acoma sag, Fig. 6).
A period of regional base-level lowering characterized bya minimum of 100 m of incision along the ancestral Rio San
Jose during the early—middle Pleistocene resulted inerosion of Q1 surfaces and formation of a deep valley 5060
m below the modern valley floor. This period of incision
lasted until ca. 0.4-0.5 m.y., at which time the Rio San Jose
valley began aggrading. A change from incision toaggradation at this time also occurred along the Rio Grandein the Albuquerque basin (Lambert, 1968), suggesting a

regional climatic control as opposed to a local tectonic influence. Episodic eruption of mid-to-late Pleistocene basalts
into the Rio San Jose and some tributary valleys obstructeddrainages and resulted in periodic upstream aggradation ofvalley floors by the streams. This damming of valleys bylava flows contributed to aggradation both along the Rio San
Jose and tributary streams during the late Pleistocene.
Valley-floor aggradation culminated after eruption of the
Laguna and Suwanee (or Cerro Verde) flows about 0.2 to 0.3
m.y. (Lipman and Mehnert, 1979; Baldridge et al., 1987;
Perry and Laughlin, unpubl. data). The timing of maximum
valley-floor aggradation within the study area appears to betime-transgressive, occurring first along downstream reaches
near the terminus of the Laguna flow (in the Laguna—Mesitaarea) and progressively later along upstream reaches (near
McCartys and Rinconada Creek). This time-transgressive end
to aggradation along the Rio San Jose is probably related to
upstream migration of the nickpoint (Fig. 5) as the Rio San
Jose cuts through the Laguna flow. The upper surface of thevalley fill is represented by the Q2 surface near Old Laguna
and by the Q3 surface between Water Creek and thecommunity of McCartys.
Renewed incision and subsequent lateral erosion of the
Rio San Jose during the late Pleistocene partially exhumed
the Laguna flow and deposited Qa13 materials as a thin
deposit capping a strath terrace near Old Laguna and along
tributaries entering the Rio San Jose downstream from Old
Laguna (Rio Paguate, Rio Moquino, and Arroyo del Valle).
Upstream from Old Laguna, the Rio San Jose continued toaggrade and tributary streams formed fan deposits gradedto the valley floor. Deposition of fine-grained Qa14 terrace
or fan deposits throughout the study area represents aperiod of aggradation during the latest Pleistocene andearly Holocene.
Valley incision followed by aggradation during the middle
to late Holocene is recorded by deposition of Qal5 materials
and burial of Qal4 deposits along the Rio San Jose. Modern

arroyos have incised 2-12 m below valley floors during thelate Holocene.
Pliocene—Quaternary deformation and uplift
Projection of surface gradients relative to present stream
profiles graded to the modern Rio San Jose valley floor
indicates the Tp1 and Tp2 surfaces are downwarped a minimum of 50 m into a broad syncline (Fig. 5). This deformation occurs in approximately the same location as the
Acoma sag (Fig. 6), a structure active in Cretaceous time(described by Kelley, 1955). This demonstrates that the
Acoma sag was active during the late Pliocene or early
Pleistocene. It is unclear when this late Cenozoic period of
deformation may have begun.
Correlative Quaternary surfaces exhibit increasing height
above base level along a southwest to northeast axis from
Laguna to Marquez and from Water Creek to Arroyo delValle (Table 4). Assuming that incision below these surfaces
is in response to uplift, then the increase in height of surfaces
above local base level indicates an increase in uplift rates
from southwest to northeast across the eastern half of thestudy area during the Quaternary. This variation in uplift
rates may be related to (1) distance from the axis of the
Acoma sag, or (2) proximity to the Rio Grande rift, withgreater uplift along the rift margin than away from the rift.
The location of erosional landforms and deposits in the
Grants—Laguna area provides evidence for active Quaternary
deformation within the Acoma sag. Qal 1 and Qal3 fan deposits are limited primarily to the area between Encinal Creek
and Rinconada Creek (Fig. 4), suggesting that Quaternary
subsidence occurred in the central part of the Acoma sag.
Erosional terrace and pediment deposits are restricted todrainages marginal to the Acoma sag (Rio Moquino, Rio
Paguate, Arroyo del Valle, and Marquez Creek) (Fig. 4).
Timing and rates of incision
Average incision rates from the late Pliocene to the present
based on the height of Tp1 above local base level range from0.07 to 0.12 m/ka (Table 5). Periods of relatively rapid
incision (approximately 0.3 to 0.7 m/ka) occurred between thetime of abandonment of Tp1 and formation of Tp2 (between
ca. 2.5 and 2.4 -± 0.18 m.y.) and between formation of Q1 andincision of the deep valley below the modern Rio San Jose
(from ca. 0.8 to 0.5 m.y.) (Table 5). The initiation of rapid
incision at ca. 2.5 m.y. corresponds to a period of rapid
extension along the margin of the Rio Grande rift from 2.6 to
1.5 m.y. , proposed by Perry et al. (1990) to accountfor an abrupt change from highly evolved rhyolite andquartzlatite magmas to less evolved latite magmas in the
Mount Taylor volcanic field. Similar incision rates
occurring from 0.8 to 0.5 m.y. may record the initiation ofa later period of extension along the rift margin.
Conclusions
Between 200 and 300 m of incision occurred in the
Grants—Laguna area during the past 2.5 m.y. in response toregional uplift. Periods of rapid incision (of as much as 0.6m/ka) occurred between ca. 2.5 and 2.4 m.y. and possibly
between 0.7-0.8 and around 0.5 m.y. Episodic entrenchment
and aggradation of fluvial systems during the late Quaternary
may be in response to uplift, climatic fluctuations, and periodic obstruction of drainages during volcanic eruptions.
Faulting was most active during the late Pliocene to early
Pleistocene. Pliocene surfaces are warped across the Acomasag, a Cretaceous structure which exhibits evidence of activity during the Pliocene and Quaternary. Tributary drainages flowing into the Acoma sag have been aggrading
during the Quaternary concurrent with formation oferosional surfaces by drainages located marginal to the sag.
Correlative Quaternary surfaces located progressively
farther from the axis of the Acoma sag exhibit an increasingheight above base level. Finally, periods of rapid uplift andIncision and more voluminous eruption of volcanic
materials between 2.5-2.4 and 0.8-0.5 m.y. may both berelated to periods of rapid extension in the transition zone
along the rift margin.
Acknowledgments
This research was completed with funding from the 1SR(Institutional Supporting Research) program at Los Alamos
National Laboratory and from the University of New Mexico. The authors would like to thank the Pueblo of Laguna
for allowing access to Pueblo lands during the course of thisinvestigation. The authors also express their appreciation to
the Cebolleta Land Grant Commission and the owners of the
L Bar Ranch for allowing access to land-grant and private
lands. Several graduate students at the University of New
Mexico, including Mark Gonzalez, Steven Harlan, andClaire Renault, volunteered help and provided valuablecriticism during the completion of field work. We also wish
to extend our thanks to Steve Reneau for a thorough manuscript review and to Anthony Garcia and James Archuleta
for their assistance in drafting the figures.

157

References
Bachman, G. 0., and Mehnert, H. H., 1978, New K-Ar dates and thelate Pliocene to Holocene geomorphic history of the central Rio
Grande region, New Mexico: Geological Society of America, Bulletin,v. 89, pp. 283-292.
Baldridge, W. S., Perry, F. V., and Shafiqullah, M., 1987, Late Cenozoic
volcanism of the southeastern Colorado Plateau: I. Volcanic geologyof the Lucero area, New Mexico: Geological Society of America,
Bulletin, v. 99, pp. 463-470.
Birkeland, P. W., 1984, Soils and geomorphology: Oxford University
Press, New York, 372 pp.
Bryan, K., and McCann, F. T., 1938, The Ceja del Rio Puerco-a borderfeature of the Basin and Range province in New Mexico, pt. II, geomorphology: Journal of Geology, v. 46, pp. 1-16.
Crumpler, L. S., 1982, Volcanism in the Mount Taylor region; in Grambling, J. A., and Wells, S. G. (eds.), Albuquerque country II: New
Mexico Geological Society, Guidebook 33, pp. 291-298.
Dethier, D. P., Harrington, C. D., and Aldrich, M. J., 1988, Late Cenozoic
rates of erosion in the western Espanola Basin, New Mexico: Evidence
from geologic dating of erosion surfaces: Geological Society of America,
Bulletin, v. 100, pp. 928-937.
Gile, L. H., Hawley, J. W., and Grossman, R. B., 1981, Soils and geomorphology in the Basin and Range area of southern New MexicoGuidebook to the Desert Project: New Mexico Bureau of Mines &Mineral Resources, Memoir 39, 222 pp.
Gordon, E. D., 1961, Geology and ground-water resources of the GrantsBluewater area, Valencia County, New Mexico: Technical Report 20,
New Mexico State Engineer, Santa Fe, 109 pp.
Grimm, J. P., 1985, The late Cenozoic geomorphic history of the Lobo
Canyon area of the Mount Taylor volcanic field, Cibola County, New
Mexico: Unpublished MS thesis, University of New Mexico, Albuquerque, 159 pp.
Harden, J. W., 1982, A quantitative index of soil development fromfield descriptions: Examples from a chronosequence in central
California: Geoderma, v. 28, pp. 1-28.
Hawley, J. W., Bachman, G. 0., and Manley, K., 1976, Quaternarystratigraphy in the Basin and Range and Great Plains provinces, New
Mexico and western Texas; in Mahaney, W. C. (ed.), Quaternary stratigraphy of North America: Dowden, Hutchinson & Ross, Inc.,
Stroudsburg, Pennsylvania, pp. 235-274.
Hunt, C. B., 1938, Igneous geology and structure of the Mount Taylor
volcanic field, New Mexico: U.S. Geological Survey, Professional
Paper 189-B, pp. 51-80.
Jenny, H., 1941, Factors of soil formation: McGraw-Hill, New York.

Kelley, V. C., 1955, Regional tectonics of the Colorado Plateau and relationship to the origin and distribution of uranium: University of New
Mexico Publications in Geology No. 5, 120 pp.
Lambert, P. W., 1968, Quaternary stratigraphy of the Albuquerquearea, New Mexico: Unpublished PhD dissertation, University of New
Mexico, Albuquerque, 300 pp.
Lipman, P. W., and Mehnert, H. H., 1979, Potassium-argon ages from
the Mount Taylor volcanic field, New Mexico: U.S. Geological Survey,
Professional Paper 1124-B, 8 pp.
Lipman, P. W., and Moench, R. H., 1972, Basalts of the Mount Taylor
volcanic field, New Mexico: Geological Society of America, Bulletin,v. 83, pp. 1335-1344.
Lipman, P. W., Pallister, J. S., and Sargent, K. A., 1979, Geologic mapof the Mount Taylor quadrangle, Valencia County, New Mexico: U.S.
Geological Survey, Map GA-1523.
Love, D. W., 1989, Geomorphic development of the Rio San Jose
valley; in Anderson, 0. J., Lucas, S. G., Love, D. W., and Cather, S.
M. (eds.), Southeastern Colorado Plateau: New Mexico GeologicalSociety, Guidebook 40, pp. 11-12.
Machette, M. N., 1985, Calcic soils of the southwestern United States;in Weide, D. L. (ed.), Soils and Quaternary geology of thesouthwestern United States: Geological Society of America, Special
Paper 203, pp. 1-22.
Maxwell, C. H., 1977, Preliminary geologic map of the McCartys quadrangle, Valencia County, New Mexico: U.S. Geological Survey, OpenFile Report 77-380.
Maxwell, C. H., 1982, El Malpais; in Grambling, J. A., andWells, S. G. (eds.), Albuquerque country II: New Mexico
Geological Society, Guidebook 33, pp. 299-301.
Moench, R. H., and Schlee, J. S., 1967, Geology and uranium depositsof the Laguna district, New Mexico: U.S. Geological Survey, Professional Paper 519, 117 pp.
Perry, F. V., Baldridge, W. S., DePaolo, D. J., and Shafiqullah, M. (inpress), Evolution of a magmatic system during continental extension:
the Mount Taylor volcanic field, New Mexico: Journal of GeophysicalResearch.
Risser, D. W., and Lyford, F. P., 1983, Water resources on the Pueblo of
Laguna, west-central New Mexico: U.S. Geological Survey, WaterResources Investigations Report 83-4038, 308 pp.
Thaden, R. E., Santos, E. S., and Raup, 0. B., 1967, Geologic map of
the Grants quadrangle, Valencia County, New Mexico: U.S.
Geological Survey, Geologic Quadrangle Map GQ-680.
Wright, H. E., Jr., 1946, Tertiary and Quaternary geology of the lower
Rio Puerco area, New Mexico: Geological Society of America, Bulletin,v. 57, pp. 383-456.

Geologic overview and Pliocene-Quaternary history of the Albuquerque basin, central New Mexico
Richard P. Lozinsky, John W. Hawley, and David W. Love
New Mexico Bureau of Mines & Mineral Resources, Socorro, New Mexico 87801

Introduction
The approximately 1000 km long Rio Grande rift forms
one of the major structural features in the southwestern
United States. The rift is composed of a series of roughly
north-south-trending, en-echelon basins (Chapin, 1971,1988) that are characterized by thin crust, elevated heat flow,
young faulting, recent volcanism, and thick basin fills(Seager and Morgan, 1978; Morgan et al. , 1986). One of thelargest and deepest basins of the rift is the Albuquerquebasin (Fig. 1). Over the past 60 years, this basin has been thefocus of many studies that have examined its stratigraphy,
structure, and complex depositional history (see previouswork section in Lozinsky, 1988). This paper presents a brief
geologic overview and discusses the Quaternary evolution ofthis important basin.
Covering an area of about 7400 km2, the Albuquerquebasin lies within the northern portion of the Mexican Highland section (Rio Grande subsection) of the Basin andRange physiographic province (Fenneman, 1931; Hawley,1986). Elevations of the margin uplifts are as much as 3050
m in the east and up to 2134 m in the west. Low
topographic relief characterizes much of the area within the
Albuquerque basin. Elevations range from about 1310 m
along the Rio Grande to around 1830 m along the eastern
piedmont slope. The valleys of the Rio Grande and Rio
Puerco are the two major erosional landforms within thebasin. A long, southward-narrowing tableland known as the
Llano de Albuquerque lies between these two valleys.
Other landforms within the basin include mesas andstructural benches, low hills and ridges, inset stream
terraces, and graded alluvial slopes.
Geologic setting
As structurally defined here, the Albuquerque basin extends from the San Felipe fault belt in the north to the Joyita

to the northwest and north, respectively. Precambrian plutonic and metamorphic rocks overlain by Paleozoic andMesozoic units crop out in the Nacimiento uplift (Woodward, 1987). The Jemez uplift is part of a large volcanic
pile and caldera complex (Smith et al., 1970; Gardner etal., 1986).
Basin structure
Although the Albuquerque basin appears topographicallyas only one basin, seismic reflection work by the Shell Oil
Company (staff comm. December 1986; Russell and Snelson,
in press) indicates that the basin is actually composed of anorthern eastward-tilted half-graben and southern westwardtilted half-graben (Fig. 2). An accommodation zone thatappears to be an extension of the Tijeras fault separates thetwo half-grabens (see Fig. 1). The half-graben morphology of
the Albuquerque basin is not unusual, most basins of the rift
display this type of morphology.
As shown in Fig. 2, the half-grabens within the Albuquerque basin are characterized by a deep, inner grabenthat is flanked by a stepped sequence of shallow, faulted
ramps that lead up to the margin areas. Faults with thelargest displacements occur several kilometers basinward
from the topographically high margins. The structural reliefof the basin is well over 10,000 m, based on the verticaldistance between the Precambrian strata on top of theeastern margin uplifts and within the deepest part of thebasin (Lozinsky, 1988).
In the north half-graben, the largest displacement fault
with as much as 2871 m of offset is located under thepresent position of the Rio Grande. Based on seismic andgroundwater data, the largest displacement fault in the south
half-graben appears to occur under the present position of
the Rio Puerco. According to seismic work by Shell Oil
Company (staff comm. December 1986; Russell andSnelson, in press), many of these faults are listric.

uplift in the south (Fig. 1). The eastern and western borderfeatures are distinctly different. The eastward-tilted SandiaManzano—Los Pinos uplift forms the prominent, relatively
linear eastern border. These uplifts consist primarily of Precambrian plutonic and metamorphic rocks which are unconformably overlain by upper Paleozoic limestone, shale,and sandstone. The western border is more irregular, morevaried in structural style, and in some places has only minor
amounts of displacement along margin faults.
The Ladron and Lucero uplifts form the southwesternbasin border. The Ladron uplift is a pyramid-shaped, structurally complex horst that consists mainly of Precambrianplutonic and metamorphic rocks with minor upper Paleozoic
units. The fault zone that separates the Lucero uplift from thebasin is extremely complex and shows Laramide-age reverseand possibly strike—slip motion with later late Cenozoic
normal faulting (Callender and Zilinski, 1976; Hammond,
1987). The gently west-dipping Lucero uplift contains upper
Paleozoic limestone, shale, and sandstone capped by lateCenozoic basalt flows.
North of the Lucero uplift, the basin boundary is marked
by the topographically subdued Puerco fault zone. This region west of the fault zone is part of the southeastern Colorado Plateau and rocks exposed here include Cretaceous
sandstone and shale, and, to a lesser extent, Jurassic clastic
units. The Nacimiento and Jemez uplifts border the basin

Basin-fill stratigraphy
The fill within the Albuquerque basin consists mostly ofsiliciclastic continental sediments with minor volcanic rocks
that can be subdivided into three units: (1) pre-Santa Fe
Tertiary deposits, (2) the Santa Fe Group, and (3) post-Santa
Fe Quaternary deposits. Pre-Santa Fe Tertiary deposits cropout only in limited exposures within the basin and weremainly studied by examining data from deep oil test wells. In
most areas of the basin, Upper Cretaceous strata underlie thebasin-fill deposits; however, along the eastern margin,
Paleozoic units underlie the basin fill deposits (Fig. 2).
Pre-Santa Fe Tertiary deposits
These deposits underlie the Santa Fe Group and indicate
the existence of at least two depositional basins that predate
the Albuquerque basin. By examining cuttings and coresfrom several of the deep oil test wells, Lozinsky (1988)
subdivided the pre-Santa Fe Tertiary deposits into the Eocene
Galisteo and Baca Formations and an overlying late Eocene
to late Oligocene unit informally referred to as the "unit of
Isleta #2 well." This informal unit generally correlates with
the Datil Group and the Oligocene volcanic sequence thatoverlies the Datil Group (Osburn and Chapin, 1983). The
Galisteo and Baca deposits consist of interbedded sandstone,siltstone, and shale derived mainly from nonvolcanic sourceareas. The Isleta #2 rocks were derived, at least partly, fromintermediate volcanic source areas and include sand-stones with locally interbedded volcanic flows and ash ofintermediate composition. Thicknesses of the Galisteo—Bacadeposits are generally less than 500 m, whereas thicknesses of
the unit of Isleta #2 range up to as much as 2185 m(Lozinsky, 1988).
Santa Fe GroupThe major basin-fill unit of the Rio Grande rift is theuppermost Oligocene to middle Pleistocene Santa Fe Group.
It is the main synrift deposit and can be subdivided into
upper and lower units based on depositional environment
and age.
The lower Santa Fe Group represents deposition in basins
with internal drainage. In the southern part of the basin, theseunits are called the Popotosa Formation by Machette (1978b,c). In the northern part, the Zia Formation of Tedford (1982)
generally correlates with the lower and middle Popotosa.
Lower Santa Fe basin fill ranges in age from 30 to 5 Ma andis dominated by intertonguing piedmont, eolian, and finegrained basin-floor deposits. The bulk of the lower Santa Fe
Group was probably deposited between 5 and 10 Ma, when
tectonism was most active in the rift.
Unlike the lower Santa Fe, the upper Santa Fe Group was
deposited in an externally drained basin. This lithostratigraphic unit is referred to as the Ceja Member by Kelley(1977) and as the Sierra Ladrones Formation by Machette(1978b, c). Upper Santa Fe Group basin fill ranges in age
from about 5 to 0.5 Ma and is characterized by intertonguing
piedmont (alluvial-fan) and basin-floor (axial-stream) deposits. Piedmont units consist of poorly sorted, weakly stratified sand and conglomerate. The basin-floor deposits include
a sand-dominated, cross-stratified axial-river facies and aninterbedded sand and clayey silt overbank facies. Abundanttrough crossbedding and high sand-to-clay ratios in the axialriver facies indicate that these sediments were deposited
within a large braided-river system. The overbank facies
represents deposition that was occurring in regions where

major braided-river systems were merging, as for example
the confluence of the ancestral Rio Grande and Rio Puerco
just west of Belen. Santa Fe Group deposition culminatedwith the construction of a broad, central-basin plain, now
partly preserved between the Rio Grande and Rio Puerco
valleys as the Llano de Albuquerque.
Several volcanic features were emplaced during the laterstages of upper Santa Fe Group deposition. Tome Hill, located about 8 km north of Los Lunas, is a small andesitic
center that stands over 100 m above the present Rio Grande
floodplain. It has been dated at 3.4 ± 0.4 Ma by Bachman
and Mehnert (1978) and is probably an exhumed feature.
Near the western edge of the Llano de Albuquerque, theundated Wind Mesa volcano forms a subdued hill along thehorizon. It is a shield volcano of probable Pliocene age thathas been considerably disrupted by faulting.
Just west of where 1-25 crosses the Rio Grande south ofAlbuquerque is the Isleta volcano, a basal maar crater overlain by several basaltic flow units. It has been dated by
Kudo et al. (1977) at 2.78 Ma. The flow units intertongue
with axial-river-facies deposits.
One of the most prominent features within the southern
Albuquerque basin is the Los Lunas volcano, located justwest of Los Lunas. This volcano consists primarily of andesite flows that have a range of K—Ar dates from 1.31 ±0.05 to 1.01 ±0.10 Ma (Bachman and Mehnert, 1978). This
volcano straddles deformed basin fill containing earlier local
andesitic eruptives and pumiceous debris flows probably
derived from the Jemez Mountains. A thick alluvial mantle
was shed from this uplifted volcano, but it is not buried bybasin fill and some of the flows descend to the east into theancestral Rio Grande valley. The locally derived mantle
contains the 1.1 Ma ash of the Upper Bandelier Tuff.
Thicknesses within the Santa Fe Group range from 1000
to 2000 m along the margin areas to as much as 4407 m inthe deeper portions of the basin. The thickest section was
penetrated in the Shell Isleta #2 test well and represents

160
the thickest documented section of the Santa Fe Group in
the entire Rio Grande rift. These thick sections are comprised mainly of the lower Santa Fe Group. Thicknesses of
the upper Santa Fe Group are generally less than 400 m.
Shell's seismic lines (staff comm. December 1986; Russell
and Snelson, in press) show that the basin fill in the south
half-graben is thickest on the west side and that it may be
as much as 4000 m thick.
Post-Santa Fe Quaternary deposits
Post-Santa Fe deposits range in age from about 1 Ma to
the present and include fan, pediment, inset terrace, volcanic, eolian, and valley-fill deposits. These units were deposited during a series of incision and partial backfilling
episodes that formed the present-day inner valley of thebasin.
Pediment and graded alluvial slopes generally occur along
the eastern side of the basin. They represent stable periods
during the incision episodes or intervals of partial valley
backfilling, and include the Airport, Llano de Manzano,
Llano de Sandia, Primero Alto, and Segundo Alto surfaces(Bryan, 1909, 1938; Bryan and McCann, 1938; Lambert,
1968; Lambert et al., 1982; Machette, 1978a, b, c, 1985).
The latter two geomorphic surfaces are related to at least
three late Pleistocene inset terrace deposits mapped along
the inner valley of the Rio Grande. They have been referred
to as the Menaul, Edith, and Los Duranes alluviums after
Lambert (1968; Lambert et al., 1982). Thicknesses of theseunits range from 10 to 40 m. The incision episodes probably
correspond to the wetter glacial periods and the backfilling
episodes to the drier interglacial periods when climate andvegetative-cover regimes were similar to the present.
Two volcanic fields were emplaced during post-Santa
Fe time. The cinder cones and basaltic flows associatedwith the Albuquerque volcanic field were emplaced
during a single short interval between 0.11 and 0.2 Ma(Geissman et al. , 1990; Bachman and Mehnert, 1978)
west of Albuquerque. The flows cover surfaces as young
as the mid-tolate Pleistocene Segundo Alto surface. The
Cat Hills volcanic field, dated at 0.13 Ma by Kudo et al.
(1977), also includes flows that spread out over the
Segundo Alto terrace (40-45 m above the river valley).
This field consists of flows and cinder cones that erupted
along a north-trending fissure.
Eolian deposits are scattered throughout the basin. The
largest eolian deposits, however, are located along the western edge of the Llano de Albuquerque and along the northside of the Rio Salado in the southern part of the basin.
The latest cut-and-fill episode of the Rio Grande is represented by the valley-fill deposits that range in age fromthe latest Pleistocene to the present. Titus (1963)
determined that these deposits were up to 36 m thick basedon water-well data in the Belen area. A valley-fill thicknessof about 24 to 36 m has been observed in all reaches of the
Rio Grande valley from the northern Albuquerque basin to
the Hueco bolson below El Paso. The valley floor has beenslowly aggrading and narrowing during most of the Holocene, because tributary-arroyo systems have beendelivering more sediment than the regional fluvial system
can remove. There is no evidence of significant valley
entrenchment and widening since the late Wisconsin, whenclimatic conditions, vegetative cover, and water andsediment discharge were considerably different (Hawley etal., 1976; Gile et al., 1981).

Pliocene and Quaternary history of the basin
The through-flowing ancestral Rio Grande drainage developed across the Albuquerque basin from the northeast by
about 5 Ma when fills of adjacent closed basins coalesced.
Two other ancestral drainages, the Rio Puerco and Rio San
Jose, joined the Rio Grande within the Albuquerque basin toform a large contributive system. The shifting channels of
this system created a large fluvial braid-plain in the eastern
third of the central basin area that probably consisted ofseveral belts of axial-stream deposits with intervening overbank deposits (Fig. 3). These fluvial belts converged southward to form a single trunk stream near Socorro. This relatively narrow trunk system, flanked by broad piedmont
alluvial plains formed by coalescent fans, continued south
through the Palomas basin south of Truth or Consequences(Lozinsky and Hawley, 1986). Distributaries of the ancestral
Rio Grande flowed into a large closed-basin area in thepresent Texas—New Mexico—northern Mexico region near
E1 Paso during the Pliocene and early Pleistocene (Gile etal., 1981). During the same period, runoff from much of the
San Luis basin watershed above Red River was blocked by
Pliocene basalts of the Taos Plateau volcanic pile as well as
local tectonic features (Heffem, 1990; Menges, 1990;
Pazzaglia and Wells, 1990).

161
Most of the deposition at this time was occurring along
the river systems in the central portion of the basin. Theseriver systems were bringing in detritus of mixed lithologies
from considerable distances outside the Albuquerquebasin. Detritus was continuing to be shed off thesurrounding uplifts, but not in the great volumes seen
during lower Santa Fe time.
Volcanism, primarily mafic, continued within the basin(i.e. emplacement of Isleta and Los Lunas volcanoes). Volcanic ashes, erupted from the Jemez caldera and as far
away as the Yellowstone area, were also deposited within
the basin (Izett, 1981).
Between about 1 and 0.5 Ma, a major change in thedepositional regime occurred when the Rio Grande fluvial
system underwent its first major incision episode, perhaps
in response to becoming integrated with drainages to theGulf of Mexico (Hawley et al., 1976; Seager et al., 1984)
and receiving increased runoff from the southern Rocky
Mountains region (Wells et al., 1987; Heffern, 1990). Thisdepositional change is a critical event in the late Cenozoic
history of the Albuquerque basin because, prior to thistime, the basin had been undergoing net aggradation.
Subsequently, however, the basin has been subjected to net
degradation as several more cycles of river incision
produced the present-day inner valley of the Rio Grande(Lambert, 1968; Hawley et al., 1976; Machette, 1985). The
initial episode of valley entrenchment also marked the endof Santa Fe Group deposition and isolated basin-floor areas(such as the Llano de Albuquerque) from further major
deposition and erosion.
Tectonism has been relatively less active in the past 0.5
Ma. Uplifts around the basin continue to shed detritus into
the basin, but not as extensively as during Santa Fe time.
There are large areas within the basin that are receiving
relatively little sediment (i.e. Llano de Albuquerque). These
areas preserve degraded fault scarps showing multiple but
relatively inconsequential movement (Bachman and Machette, 1977). Significant aggradation in the major drainages
during back-filling episodes is documented by thick, inset
terrace deposits and undissected valley fill (Lambert, 1968;
Hawley et al., 1976). Major fluvial systems are still delivering mixed lithologies derived in part from outside thebasin. Volcanism also continued as the Albuquerque and Cat
Hills volcanic fields developed on the Llano de Albuquerque.
Following a major interval of fluvial incision to levels 25 to
40 m below modern valley floors, probably between 25,000
and 10,000 yrs B.P., the Rio Grande base level has beenslowly rising, accompanied by deposition of latest
Quaternary valley fill (Hawley et al., 1976).
Acknowledgments
The work on the Santa Fe Group was part of Lozinsky's
PhD dissertation completed at the New Mexico Institute of
Mining & Technology. Funding was provided by the New
Mexico Bureau of Mines & Mineral Resources (Frank E.
Kottlowski, director) and by the New Mexico GeologicalSociety.
References
Anonymous, 1961, Structural problems of the Rio Grande trough in Albuquerque country: New Mexico Geological Society, Guidebook 12,
pp. 144-147.
Bachman, G. 0., and Machette, M. N., 1977, Calcic soils and calcretes inthe southwestern United States: U.S. Geological Survey, Open-file
Report 77-794, 163 pp.

Bachman, G. 0., and Mehnert, H. H., 1978, New K–Ar dates and the late
Pliocene to Holocene geomorphic history of the central Rio Grande
region, New Mexico: Geological Society of America, Bulletin, v. 89,
pp. 283-292.
Bryan, K., 1909, Geology of the vicinity of Albuquerque: University of
New Mexico Bulletin 51, Geological Series, v. 3, 24 pp.
Bryan, K., 1938, Geology and groundwater conditions of the Rio Grandedepression in Colorado and New Mexico; in U.S. Natural Resources
Planning Board, The Rio Grande Joint Investigations in the upper Rio
Grande Basin: U.S. Government Printing Office, Washington, D.C., v.
1, pt. 2, pp. 197-225.
Bryan, K., and McCann, F. T., 1937, The Ceja del Rio Puerco: A borderfeature of the Basin and Range province in New Mexico, Part I: Journalof Geology, v. 45, pp. 801-828.
Bryan, K., and McCann, F. T., 1938, The Ceja del Rio Puerco: A borderfeature of the Basin and Range province in New Mexico, Part II:
Journal of Geology, v. 46, pp. 1-16.
Callender, J. C., and Zilinski, R. E., 1976, Kinematics of Tertiary and
Quaternary deformation along the eastern edge of the Lucero uplift,
central New Mexico: New Mexico Geological Society, Special Publication no. 6, pp. 53-61.
Chapin, C. E., 1971, The Rio Grande rift, Part I—modifications andadditions: New Mexico Geological Society, Guidebook 22, pp. 191201.
Chapin, C. E., 1988, Axial basins of the northern and central Rio Grande
rifts; in Sloss, L. L. (ed.), Sedimentary cover—North American Craton(U.S.): Geological Society of America, Geology of North America, v.
D-2, pp. 165-170.
Fenneman, M. N., 1931, Physiography of the western United States:
McGraw–Hill, New York, 534 pp.
Gardner, J. N., Goff, F., Garcia, S., and Hagan, R. C., 1986, Stratigraphic
relations and lithologic variations in the Jemez volcanic field, New
Mexico: Journal of Geophysical Research, v. 91, pp. 1763-1778.
Geissman, J. W., Brown, L., Turrin, B. D., McFadden, L. D., and Harlan,
S. S., 1990, Brunhes chron excursion/polarity episode recorded during
the late Pleistocene, Albuquerque Volcanoes, New Mexico, U.S.A.:
Geophysical Journal International, v. 102, pp. 73-88.
Gile, L. H., Hawley, J. W., and Grossman, R. B., 1981, Soils and geomorphology in the Basin and Range area of southern New Mexico—
Guidebook to the Desert Project: New Mexico Bureau of Mines &Mineral Resources, Memoir 39, 222 pp.
Hammond, C. M., 1987, Geology of the Navajo Gap area between the
Ladron Mountains and Mesa Sarca, Socorro County, New Mexico: astructural analysis: Unpublished MS thesis, New Mexico Institute of
Mining & Technology, Socorro, 212 pp.
Hawley, J. W., 1986, Physiographic provinces; in Williams, J. L. (ed).,
New Mexico in maps (2nd edition): University of New Mexico Press,
Albuquerque, pp. 23-27.
Hawley, J. W., Bachman, G. 0., and Manley, K., 1976, Quatemarystratigraphy in the Basin and Range and Great Plains provinces, New
Mexico and western Texas; in Mahaney, W. C. (ed.), Quaternary stratigraphy of North America: Dowden, Hutchinson, and Ross, Inc.,
Stroudsburg, Pennsylvania, pp. 235-274.
Heffern, E. L., 1990, A geologic overview of the Wild Rivers Recreation
Area, New Mexico: New Mexico Geological Society, Guidebook 41,
pp. 229-236.
Izett, G. A., 1981, Volcanic beds: recorders of upper Cenozoic silicic
pyroclastic volcanism in the western United States: Journal of Geophysical Research, v. 86, pp. 10200-10222.
Kelley, V. C., 1977, Geology of Albuquerque basin, New Mexico: New
Mexico Bureau of Mines & Mineral Resources, Memoir 33, 59 pp.
Kudo, A. M., Kelley, V. C., Damon, P. E., and Shafiqullah, M., 1977, K–
Ar ages of basalt flows at Canjilon Hill, Isleta volcano, and the Cat
Hills volcanic field, Albuquerque–Belen basin, central New Mexico:
Isochron/West, no. 18, pp. 15-16.
Lambert, P. W., 1968, Quaternary stratigraphy of the Albuquerque area,
New Mexico: Unpublished PhD dissertation, University of New
Mexico, Albuquerque, 257 pp.
Lambert, P. W., Hawley, J. W., and Wells, S. G., 1982, Supplemental
road-log segment III-S: Urban and environmental geology of the Albuquerque area: New Mexico Geological Society, Guidebook 33, pp.
97-119.
Lozinsky, R. P., 1988, Stratigraphy, sedimentology, and sand petrology of
the Santa Fe Group and pre-Santa Fe Tertiary deposits in the Albuquerque basin, central New Mexico: Unpublished PhD dissertation,
New Mexico Institute of Mining & Technology, Socorro, 298 pp.
Lozinsky, R. P., and Hawley, J. W., 1986, The Palomas Formation of

162
south-central New Mexico-a formal definition: New Mexico Geology,
v. 8, no. 4, pp. 73-78, 82.
Machette, M. N., 1978a, Dating of Quaternary faults in the southwestern
United States by using buried calcic paleosols: U.S. Geological Survey,
Journal of Research, v. 6, pp. 369-381.
Machette, M. N., 1978b, Geologic map of the San Acacia quadrangle,
Socorro County, New Mexico: U.S. Geological Survey, Geologic
Quadrangle Map GQ- 1415, scale 1:24,000.
Machette, M. N., 1978c, Preliminary geologic map of the Socorro 1° x 2°
quadrangle, central New Mexico: U.S. Geological Survey, Open-file
Report 78-607, scale 1:250,000.
Machette, M. N., 1985, Calcic soils of the southwestern United States:
Geological Society of America, Special Paper 203, pp. 1-21.
Menges, C. M., 1990, Late Cenozoic rift tectonics and mountain-front
landforms of the Sangre de Cristo Mountains near Taos, northern New
Mexico: New Mexico Geological Society, Guidebook 41, pp. 113-122.
Morgan, P., Seager, W. R., and Golombek, M. P., 1986, Cenozoic
thermal, mechanical and tectonic evolution of the Rio Grande rift:
Journal of Geophysical Research, v. 91, pp. 6263-6276.
Pazzaglia, F. J., and Wells, S. G., 1990, Quaternary stratigraphy, soils andgeomorphology of the northern Rio Grande rift: New Mexico Geological Society, Guidebook 41, pp. 423-430.
Osburn, G. R., and Chapin, C. E., 1983, Nomenclature for Cenozoic rocksof northeast Mogollon-Datil volcanic field, New Mexico: New Mexico
Bureau of Mines & Mineral Resources, Stratigraphic Chart 1.
Russell, L. R., and Snelson, S. (in press), Structural style and tectonic
evolution of the Albuquerque basin segment of the Rio Grande rift; in
Pinet, B., and Bois, C. (eds.), Potential of deep seismic profiling forhydrocarbon exploration: French Petroleum Institute, Research Conference Proceedings.
Seager, W. R., and Morgan, P., 1978, Rio Grande rift in southern New

Mexico, west Texas, and northern Chihuahua; in Riecker, R. E. (ed.),Rio Grande rift: tectonics and magmatism: American GeophysicalUnion, Washington, D.C., pp. 87-106.
Seager, W. R., Shafiqullah, M., Hawley, J. W., and Marvin, R. F., 1984,
New K-Ar dates from basalts and the evolution of the southem Rio
Grande rift: Geological Society of America, Bulletin, v. 95, pp. 87-99.
Smith, R. L., Bailey, R. A., and Ross, C. S., 1970, Geologic map of theJemez Mountains, New Mexico: U.S. Geological Survey,
Miscellaneous Investigations Series, Map 1-571, scale 1:250,000.
Tedford, R. H., 1982, Neogene stratigraphy of the northwestern Albuquerque basin: New Mexico Geological Society, Guidebook 33, pp.
273-278.
Titus, F. B., 1963, Geology and ground-water conditions in eastern Valencia County, New Mexico: New Mexico Bureau of Mines & Mineral
Resources, Ground-water Report 7, 113 pp.
Wells, S. G., Kelson, K. I., and Menges, C. M., 1987, Quaternary evolution of fluvial systems in the northern Rio Grande rift, New Mexico
and Colorado: implications for entrenchment and integration ofdrainage systems; in Menges, C. M. (ed.), Quatemary tectonics,
landform evolution, soil chronologies and glacial deposits-northern Rio
Grande rift of New Mexico: Friends of the Pleistocene Field Trip
Guidebook, pp. 55-69.
Woodward, L. A., 1987, Geology and mineral resources of Sierra Nacimiento and vicinity, New Mexico: New Mexico Bureau of Mines &Mineral Resources, Memoir 42, 84 pp.
Woodward, L. A., Callender, J. F., Seager, W. R., Gries, J. C., Shaffer, W.
L., and Zilinski, R. E., 1978, Tectonic map of the Rio Grande rift regionin New Mexico, Chihuahua, and Texas; in Hawley, J. W. (compiler),
Guidebook to the Rio Grande in New Mexico and Colorado: New
Mexico Bureau of Mines & Mineral Resources, Circular 163, Sheet 2.

Cuspate-lobate folds along a sedimentary contact, Los Lunas volcano, New Mexico
William C. Haneberg
New Mexico Bureau of Mines & Mineral Resources, Socorro, New Mexico 87801

BackgroundCuspate-lobate folds are characteristic of strongly compressed interfaces separating media of differing mechanicalcompetence, with the lobate folds cored by the more competent material and the cuspate folds cored by the less competent material (e.g. Ramsay and Huber, 1987, p. 403). Other
commonly used terms for this type of feature include
mullion, fold mullion, and arc-and-cusp structure. Althoughcuspate-lobate folds are commonly associated with highgrade metamorphic rocks, their morphology is similar to thatof alternating load casts and flame structures in sedimentary
sequences. In addition to having geologic significance as
indicators of both competence contrast and shortening
direction, cuspate-lobate folds are also amenable to theoretical and experimental analysis.
Ramsay and Huber (1987, pp. 392-393) maintain that
cuspate-lobate folds form only along interfaces separating
media of low viscosity contrast, perhaps less than an orderof magnitude. Smith (1975, 1977) and Fletcher (1982) analyzed the growth of instabilities along interfaces separating
both linear and nonlinear fluids, and concluded that, although so-called mullion instabilities have small growth
rates, mullion growth is enhanced in nonlinear fluids. One
important conclusion of such analyses is that no first-order
dynamic instability, which would maximize the growth ofstructures with a preferred wavelength, will develop between two semi-infinite fluids. Instead, there is only kinematic amplification of pre-existing perturbations along theinterface. Johnson and Pfaff (1989) provided a detailed geometric analysis of the waveforms necessary to produce different forms of folds, including cuspate-lobate folds, andshowed how cuspate-lobate folds might evolve from sinusoidal folds in linear viscous multilayers subjected to shortening of 6-8%. Recent experimental work with strainsoftening silicon putty models (Sokoutis, 1990) has shownthat cuspate-lobate structures grow rapidly along a single
interface separating two semi-infinite fluids with low viscosity contrast, but again with no preferred wavelength,
subjected to shortening greater than 10%.
Observations
A series of exceptionally well-developed cuspate-lobate
folds is exposed along a sedimentary contact at Los Lunas
volcano (Fig. 1). Below the contact is a buff to orange
alluvium composed of silt and fine sand, with no distinct
bedding. Above the contact are a 1-2 cm thick white ash, a
10-15 cm thick gray tephra ranging in size from sand to
granule, and a coarser orange tephra up to 1 m thick. Muchof the less-resistant alluvium beneath the contact has been
removed by erosion, showing that both cusps and lobes
extend at least a meter back into the outcrop; therefore, thesefolds can be analyzed as two-dimensional structures. Most
cusps in the sequence point downward, suggesting that the
finer-grained alluvium was generally the more competentmaterial during deformation; however, a few very lowamplitude cusps do point upward. The lobe portions of

the folds are somewhat flat, similar to a theoretical patternconsisting of a positive first waveform and a negative second
waveform (Johnson and Pfaff, 1989, p. 128). Cusp amplitude
and peakedness both decrease along the ash and tephra
contacts above the alluvium. In addition to the prominentcuspate—lobate folds, the interface is also folded into a gentle
syncline with a wavelength on the order of 102 m, and cut by
two small thrust faults with shortening of 0.19 and 0.48 m.
The distribution of 54 fold wavelengths, measured bystretching a tape between adjacent cusps, has a fairly strong
central tendency (Fig. 2), with an arithmetic mean of 1.78 m
and a standard deviation of 0.76 m.
Conclusions
The existence of both cuspate—lobate folds and small thrustfaults shows that the sedimentary contact described in thispaper has been shortened significantly, and cusp orientationshows that the alluvium was from one to ten times morecompetent than the volcaniclastic sediments during deformation. If mass was conserved during shortening, the thickness of the stratigraphic section must also have grownproportionally. Although measurements of arc length are
more appropriate for preferred wavelength studies (Sherwin
and Chapple, 1968), straight-line cusp-to-cusp measurements
are much easier to collect and provide a useful first
approximation of wavelength distribution. Because neither
theory nor experiment predict the existence of a preferred
wavelength for cuspate—lobate folds along a single interface,
we are faced with two alternative explanations for our observation of somewhat uniform wavelength. First, the cuspate—lobate folds could represent the amplification of preexisting perturbations along the interface, for example ripple
marks. Second, it is possible that a preferred wavelength

evolved because the sediments on one or both sides of theinterface behaved as finite layers, for which dynamic instabilities will arise. Smith (1975), for example, suggests thatmany cuspate—lobate folds along single interfaces in thefield are actually the erosional remnants of a finite layer withcuspate—lobate folds along both interfaces. At present, meager knowledge of the depositional and deformational historyof these sediments does not allow either of these explanations
to be favored over the other. It is hoped, however, thatdetailed mapping and mechanical analysis will provide a
better understanding of the exceptional structures at Los
Lunas volcano.
Acknowledgments
This work was supported by the New Mexico Bureau of
Mines & Mineral Resources. David Love introduced me toa number of structural problems along the flanks of Los
Lunas volcano, assisted with wavelength measurements,
and, along with John Hawley, read drafts of this contribution.
References
Fletcher, R. C., 1982, Analysis of the flow in layered fluids at small, but
finite, amplitudes with application to mullion structures:
Tectonophysics, v. 81, pp. 51-66.
Johnson, A. M., and Pfaff, V. J., 1989, Parallel, similar and constrained
folds; in Johnson, A. M., Burnham, C. W., Allen, C. R., and Muehlberger, W. (eds.), Richard H. Jahns Memorial Volume: Engineering
Geology, v. 27, pp. 115-180.
Ramsay, J. G., and Huber, M. I., 1987, The techniques of modem structural geology, Vol. 2: Folds and fractures: Academic Press, London,
700 pp.
Sherwin, J., and Chapple, W. M., 1968, Wavelengths of single layerfolds: a comparison between theory and observation: American Journalof Science, v. 266, pp. 167-179.
Smith, R. B., 1975, Unified theory of the onset of folding, boudinage, andmullion structure: Geological Society of America, Bulletin, v. 86, pp.
1601-1609.
Smith, R. B., 1977, Formation of folds, boudinage, and mullions in nonNewtonian materials: Geological Society of America, Bulletin, v. 88,
pp. 312-320.
Sokoutis, D., 1990, Experimental mullions at single and doubleinterfaces: Journal of Structural Geology, v. 12, pp. 365-373.

164
Hydrogeologic investigations of upper Sierra Ladrones Formation
Ruth C. Lohmann, J. Matthew Davis, David W. Love, and Fred Phillips
New Mexico Institute of Mining & Technology, Socorro, New Mexico 87801

The Sierra Ladrones Formation (Machette, 1978a, b) is theupper unit of the Santa Fe Group, which comprises the basin
fill of the Albuquerque basin (seen at tour Stop 2-4).
Hydrogeologic mapping of an upper Sierra Ladrones section
near Bosque, New Mexico, is currently being conducted by
the Hydrology Division of the New Mexico Tech Geoscience
Department. The goal of this research is to study the relationships between the observed pattern of geologic facies, thedepositional processes which formed these facies, and theresulting patterns of permeability. Very detailed geologic
mapping of the exposed deposits in conjunction with intensive sampling of the permeability of the deposits form thebasis of the research effort.
The upper Sierra Ladrones Formation is composed of an
alluvial—fluvial—eolian sequence of sediments. Mappingof the deposits has been focused on grouping the observed
geologic facies into architectural elements as defined by
Miall (1985). Miall defines an architectural element as a"three-dimensional [rock or sedimentary] body characterized by distinctive lithofacies assemblages, external geometry, and orientation . . . the most distinct characteristicof . . . [which] is that it consists of genetically related lithofacies. . . ."Three of the fluvial architectural elements defined by Miall
have been recognized and mapped at the southern field site.
Fig. 1 is a schematic section of the architectural elements
delineated at the site. As shown, there are two different
channel elements, CH-I being a higher-energy element dominated by medium to coarse sands and gravels, and CH-II
which is characterized by fine to medium sands with abundant interbedded clay and silt. The overbank fines element
OF is composed of interbedded silts and clays, as well as
thin, immature, dominantly clay paleosols. The fourth element recognized at the site is the paleosol element P. Paleosols at the field site are both sandy and clay-rich,indicating differing parent materials. The paleosols are lessmature than the soil on the Llano de Albuquerque, exhibiting
only stage I to II of the carbonate morphogenetic sequenceof Gile et al. (1966). Miall includes the paleosol facies in the
overbank fines element due to the lack of a fourth-orderbounding surface at their base. However, due to their uniquepermeability characteristics, the overbank fine element of
Miall (1985) is split into a mappable paleosol element andthe traditional overbank fine element.
The character and spatial arrangement of these mapped
elements allow additional interpretations of the processes
which formed the upper Sierra Ladrones Formation. First,
the presence of two distinctly different channel elementssuggests that at least two types of stream systems may havelaid down these deposits. The higher-energy CH-I element
resembles a fluvial deposit of a perennial stream, possibly
the ancestral Rio Grande. On the other hand, the lowerenergy CH-II elements commonly have characteristics ofephemeral-stream alluvium (abundance of clay drapes,
sand-size clay clasts, and armored mud balls) and are on theorder of a few tens of meters wide, similar to the scale ofthe modern Rio Puerco.
The vertical succession of the elements suggests that thedeposits were laid down by rivers which avulsed laterally insmall and large steps. Using the alluvial models of Allen165(1974) as a guide, several sedimentary cycles of river movement are visible in Fig. I. The base and top of these cycles
are marked by channel elements, with a vertical successionof overbank and paleosol elements in between. If an overbank fine element directly overlies the basal channel element, it indicates that the channel avulsed away from its
original position in a small step, leaving its floodplain overlying the channel deposit. An overlying paleosol followedby another overbank and then the top channel element implies that the channel moved yet farther away and thenreturned in small steps to its original location over aninterval as long as thousands of years. A second type ofcycle seen in the section is one in which a paleosol element
directly overlies the basal channel element. This suggeststhat the channel moved away in a larger avulsive step,leaving its original depositional site starved of sediment.
Both such patterns are seen in the section, suggesting that
small and large avulsive river movements were responsiblefor deposition of the observed geologic-facies assemblages.
Another interpretation of the depositional processes active inthe formation of these sediments stems from the paleosol
models of Kraus and Bown (1987). According to their models,overall vertical changes in paleosol maturity within an alluvial
sequence indicate overall changes in the rate of sedimentation.
In a measured section that includes the mapped portion of theupper Sierra Ladrones Formation, paleosol maturity is greatest
midway through the section and decreases markedly upwards(except for the soil developed on

the Llano de Albuquerque summit; see Machette, 1985). In
fact, vertically above the mapped section the frequency ofthe paleosols and overbank fines decreases notably, withpaleosols becoming very thin and immature. In generalterms, this upward decrease in the presence and maturity of
paleosols suggests that there was an overall increase in therate of sedimentation or in the rate of lateral reworking.
Possible causes of such an increase include a shift to awetter climate and/or a cessation of basin subsidence. The
latter cause would result in decreased preservation of soils
and overbank deposits as they would be reworked by rivers
shifting across the basin floor.
Assessment of the permeability patterns within this fluvial—interfluvial sequence is achieved with a portable airpermeameter. The air permeameter measures the steadystate air-flow rate for a prescribed pressure. The air-flowrate data can then be converted to permeability according
to the work by Goggin et al. (1988). Two small-scalestudies of permeability patterns have been performed, one
within the lowermost CH-I1 element and the other within
the CHI element. Study of permeability at the larger scaleis currently ongoing.
Fig. 2 illustrates the results of the CH-II study. In thisstudy, 277 air-flow-rate measurements were collected froma 2 m wide by 0.5 m high outcrop. Fig. 2a illustrates thegeologic facies and sampling locations. The histogram ofthe data (Fig. 2b) indicates a log-normal distribution of
permeability. Fig. 2c shows the horizontal variogram esti-

166
mated with the log data. As shown, an exponential modelwith a correlation length of approximately 40 cm wasfitted to the data.
The results of the CH-I study, not shown here (seeDavis, 1990), indicate that the permeability data are also
log-normally distributed. In this case, a nested exponential
model was fitted to the horizontal variogram. The nested
structure is attributed to the two different lithologies present
in the CH-I outcrop.
Our preliminary analysis suggests that the upper part of the
Sierra Ladrones exposed along the Llano de Albuquerqueescarpment is the product of deposition of the ancestral Rio
Grande and its tributaries to the west. The patterns of permeability appear to coincide well with the mappable geologic
features. Ultimately, we hope to develop a means of quantifying patterns of permeability using sedimentologically
based models.
This work is being funded by the U.S. Department ofEnergy, Subsurface Science Program.
References
Allen, J. R. L., 1974, Studies in fluviatile sedimentation: Implications forpedogenic carbonates units, Lower Old Red Sandstone, Anglo-Welsh
outcrop: Geological Journal, v. 9, pp. 181-208.
Davis, J. M., 1990, An approach for the characterization of spatial variability in the Sierra Ladrones Formation, Albuquerque Basin, central
New Mexico: Unpublished MS thesis, New Mexico Institute of Mining& Technology, Socorro, 137 pp.
Gile, L. H., Peterson, F. F., and Grossman, R. B., 1966, Morphologicaland genetic sequences of carbonate accumulation in desert soils: Soil
Science, v. 101, pp. 347-360.
Goggin, D. J., Thrasher, R., and Lake, L. W., 1988, A theoretical andexperimental analysis of minipermeameter response including gas slippage and high velocity flow effects: In Situ, v. 12, pp. 79-116.
Kraus, M. J., and Bown, T. M., 1986, Paleosols and time resolution in
alluvial stratigraphy; Chapter 6 in Wright, V. P. (ed.), Paleosols: Their
recognition and interpretation: Princeton University Press, Princeton,
New Jersey, pp. 180-207.
Machette, M. N., 1978a, Geologic map of the San Acacia quadrangle,
Socorro County, New Mexico: U.S. Geological Survey, Geologic
Quadrangle Map GQ-1415, scale 1:24,000.
Machette, M. N., 1978b, Preliminary geologic map of the Socorro 1°x 2°
quadrangle, central New Mexico: U.S. Geological Survey, Open-file
Report 78-607.
Machette, M. N., 1985, Calcic soils of the southwestern United States:
Geological Society of America, Special Paper 203, pp. 1-21.
Miall, A. D., 1985, Architectural-element analysis: A new method of
facies analysis applied to fluvial deposits: Earth Science Reviews, v.
22, pp. 261-308.

Effect of climatic change on Estancia Valley, New Mexico: Sedimentation and landscape evolution
in a closed-drainage basin
Bruce D. Allen
Department of Geology, University of New Mexico, Albuquerque, New Mexico 87131

Introduction: Estancia Valley as aclimatic—hydrologic model
For over 80 years, Estancia Valley in central New Mexico
has attracted the attention of geologists and hydrologists
interested in the basin's natural history. The first comprehensive study of the basin was conducted by 0. E. Meinzer
during the summer of 1909. Meinzer's primary task was toreport on the feasibility of ground-water irrigation in Estancia
Valley. He also assisted in the classification of land under thehomestead act. Meinzer (1911), in just six weeks of field
work, produced a classic paper which provided a conceptualmodel of the interplay between geology, climate, andhydrology. Meinzer's paper also contained an extensive
description of the geomorphic and stratigraphic evidence foran ancient lake in Estancia Valley.
Subsequently, Antevs (1935, 1954) and Leopold (1951)
refined the evidence for past climatic changes in Estancia
Valley and formulated relationships between specific climatic variables and the hydrologic balance in the topographically closed basin. Their objective was to derive estimates ofglacial-age climate for this region by using semi-quantitativemodels that were constrained by the evaporative surface area
represented by the preserved shorelines of ancient Lake
Estancia. Estancia Valley was well suited for this exercise
because the absence of mountain glaciers in the Estancia
drainage basin during the late Pleistocene eliminated acomplicating factor (the storage of precipitation as ice and its
delayed release as meltwater) that was present in many of theGreat Basin "pluvial" lake basins. Antevs

and Leopold concluded that climate during the last glacial
maximum was both wetter and colder than today.
Since these pioneering efforts, several researchers havederived other estimates of paleoclimate for the Estancia
area by using different initial assumptions (see review inSmith and Anderson, 1982). The intent of this paper is not
to determine which of these climatic estimates is mostcorrect, but to provide a brief description of the geologic
evidence for hydrologic—climatic change in the area forthe time interval from the last glacial maximum to thepresent. The discussion of the late Pleistocene lacustrine
stratigraphic record is based on the author's ongoing
research. Geomorphic records in the Estancia Valley
provide the primary evidence for changes in the Holocene.
Hydrogeologic setting
Estancia Valley (Fig. I) lies in the Sacramento section of
the Basin and Range physiographic province, which is transitional to the Great Plains physiographic province to the east
(Hawley, 1986). Playa lakes currently occupy an extensive
complex of nearly 100 deflation basins that are cut into thevalley floor in the southern part of the basin. Topographic
closure between the valley floor (1852 m) and the sill with
the Pinos Wells basin to the southeast is about 80 m. The
Manzano Mountains, which form the western margin of thebasin, exhibit the highest elevations, with peaks of over 3050
m. Paleozoic marine and terrestrial strata dip gently eastward
into the Estancia Valley from the Manzano Mountains andoccur at depth under most of the basin (Smith,167
lost from the Estancia basin through shallow subsurface
flow, although it is reasonable to assume that some leakage
and deep interbasin flow occurs through the Paleozoic carbonate strata underlying the basin (Jenkins, 1982).
DeBrine (1971) monitored hydrologic and climaticconditions at one of the Estancia playas for two years andestimated that 33 to 45 million m3 of ground water is
evaporatively discharged from the playas each year,
assuming negligible subsurface leakage.
A substantial contribution to the volume of Lake
Estancia in the late Pleistocene must have been provided bygroundwater discharge. The area of the drainage basin, andespecially the recharge area, was quite small with respect tothe area of the lake (Fig. I). The combination of continuous
ground-water discharge and a large evaporative surface area
produced a lake which rose and fell dramatically during thelate Pleistocene, but not to the point of complete
desiccation and deflation. As a result, sedimentation in theancient lake basin was essentially uninterrupted andcombined hydrologic and climatic sensitivity with acontinuous stratigraphic record.
Stratigraphic and sedimentational expressionof climatic change1957). Precambrian crystalline rocks of the Pedernal Hills
are exposed on the eastern drainage divide, where
elevations are less than 2285 m.
Most geologists working in the area, beginning with Darton (1928), surmised the existence of a north—south-trending
fault zone that downdrops Estancia Valley along its eastern
margin. Smith (1957) and subsequent workers, citing theapparent absence of Paleogene strata in the valley fill, havesuggested that formation of the modern basin was related tolate stages in the development of the Rio Grande rift.
Smith (1957) divided the sequence of unconsolidated
sediments in Estancia Valley into the alluvial "valley fill,"
which he considered to be of late Pliocene(?) and Pleistocene age, and younger "lake and dune deposits." The combined thickness of the basin-filling sediments is over 122
m in the axial portion of the basin (New Mexico StateEngineer Office, 1967), becoming thinner towards themargins and overlapping Paleozoic bedrock at an elevationof approximately 2010 m along the eastern slope of the
Manzano Mountains (Fig. 1).
The sediments of the alluvial valley fill form a large aquiferthat is utilized for irrigation on the western side of the basin.
Recharge to the aquifer is derived principally fromprecipitation at higher elevations, especially in the Manzano
Mountains (DeBrine, 1971). A map of the total dissolved
solids in the Estancia ground-water basin (New Mexico
Engineer Office, 1965) reveals a pattern of concentric contour lines that converge along the axis of the basin, where
several of the modern playas seasonally precipitate halite.
The pattern of increasing solute load towards the centerof the basin suggests that little ground water is presently

Lacustrine sediments
Up to 10 m of the lacustrine sequence is exposed in thewalls of the numerous deflation basins in the southern partof the basin (Fig. 1). Examination of water-well cuttings
and well logs from the southern part of the basin by Titus(1969) suggests that the lacustrine sequence may be morethan 30 m thick, with an interbedded quartz-sand interval("medial sand") occurring at a depth of about 18 m. The
discussion here is limited to the upper part of the exposedlacustrine sequence, which includes the interval deposited
during the last glacial maximum. Bachhuber (1989) discusses the stratigraphy of the entire exposed sequence.
Sediments deposited in Lake Estancia consist primarilyof carbonate-rich clay and gypsum silt and sand. The larger
gypsum grains are morphologically similar to "type II" and"type IV" lenticular gypsum crystals grown by Cody and
Cody (1988) under analog saline terrestrial conditions, andcommonly show evidence of abrasion. The size, shape, andmechanical weathering of the gypsum grains suggest thatthey grew in muddy areas around the margin of the lake
where they were exposed, sorted, and eventually carried
into the lake. A marginal gypsum source area implies thatstratigraphic variations in gypsum content reflect past fluctuations in: (1) the areal extent of the mudflats and theirproximity to the basin center (a function of lake area), and(2) the amount of wind and wave energy available for transport from marginal areas to the basin center (a function ofwind intensity). The proportion of gypsum, carbonates, andclays in the lake sediments, whose distribution was controlled by the rise and fall of the lake and fluctuations inwind intensity, underwent significant change that can berecognized at all time scales.
Evidence for millennial-scale climatic cycles
Two major sedimentary facies can be recognized in the
exposed lake sequence, a laminated facies and a non-laminated (bioturbated) facies. The two facies alternate vertically on a scale of decimeters to meters. The absence ofdelicate sedimentary structures in the non-laminated facies

168
is due to burrowing activity, probably by aquatic oligochaete
worms, as indicated by the size and shape of infilled burrows.
Most aquatic oligochaetes are able to thrive in low
concentrations of dissolved oxygen (Pennack, 1989), suggesting that changes in salinity and presumably lake level,
rather than disaerobic conditions, were principally responsible
for the elimination and introduction of bioturbating organisms.
The presence of fossil trout (Oncorhynchus clarki) and freshwater mollusks and ostracodes in bioturbated intervals(Bachhuber, 1971, 1989) supports the hypothesis that thebioturbated facies developed with freshening of the lake.
Decimeter- and meter-scale alternations between laminatedand bioturbated sediments are therefore thought to reflectchanges in the balance between inflow and evaporation, water
depth, and surface area of Lake Estancia.
The upper 5 m of the lacustrine sequence consist of two
such bioturbated intervals and two laminated intervals(Fig. 2), which can be correlated across the central part ofthe lake basin. Radiocarbon dates from the lacustrine
sequence (V. Markgraf, pers. comm. 1988; Bachhuber,
1989; this study) indicate that these major facies changes
reflect fluctuations in lake volume and area that occurred
over 2 to 4 kyrs.
Three radiocarbon dates obtained from bulk sediment
samples were not used in establishing a chronostratigraphy
for the Estancia sequence. Microscopic examination of thebulk material represented by unaccepted radiocarbon dates"A" and "C" (Fig. 2; 9830 ± 240 and 13,450 ± 170 yrs B.P.,
respectively) revealed the presence of rootlet tubes, suggesting that contamination by modern organic material may
be reflected in the dates obtained from these samples. Another sample, collected from sediments deposited during asaline phase near the top of a laminated interval, yielded aradiocarbon date of 15,410 ± 650 yrs B.P. Thompson et al.
(1990) note anomalously old radiocarbon dates from dispersed organic material deposited during saline phases oflakes in the Great Basin, and a similar association may be
responsible for what appears to be an anomalously old date
for the Estancia material (date "B" in Fig. 2).
The lower bioturbated interval (Fig. 2) was deposited
during a highstand of the lake that was contemporaneouswith the last glacial maximum. A 2 m sequence oflaminated sediments above the glacial-maximum sediments
indicates a substantial decrease in the size of Lake Estancia
about 15,000 yrs B.P. A progressive lowering of lake level
during the deposition of the laminated interval near 15,000
yrs B.P. (Fig. 2) is indicated by the upward appearance of
sedimentary structures associated with wave/bottom
interactions. The upper bioturbated interval represents
another highstand near 13,000 yrs B.P. Bioturbatedsediments deposited during the last highstand are in turnoverlain by laminated sediments that signify a decline inlake area and, ultimately, desiccation of the lake. The
timing of the last two major highstands in Estancia Valley
is similar to the record of highstands in other dated paleolakes near this latitude in the western U.S. (Allen andAnderson, 1989), including paleo-lakes Mojave, California(Wells et al., 1989), and Cochise, Arizona (Waters, 1989).
Evidence for centennial-scale climatic changes
Centennial changes in lake area and volume are recordedin the stratigraphic section as relative changes in carbonates,
clay, and gypsum silt and sand. Relationships between relative changes in carbonates and clay are not yet clear, andthe discussion of centennial and decadal changes in this andthe following section is limited to gypsum.
Variations in the flux of gypsum during highstands of thelake are represented in the basin center as stratigraphic
changes in relative percent of gypsum (Fig. 3). Time-series
based on analyses at 2.5 cm contiguous intervals, and using
average accumulation rates to extrapolate between
radiocarbon dates, reveals significant variability in gypsum
flux during highstands of the lake. In the portion of theglacial-maximum sequence deposited between —18,000
and 15,800 yrs B.P., five major episodes of gypsum
transport can be recognized, with an average period of about
560 years (Fig. 3A).
Major changes in transport of gypsum to the center of thebasin were probably accompanied by contraction andexpansion in the area of marginal mudflats and a corresponding rise and fall of lake level. As lake levels rose, anincreasing proportion of the marginal mudflats that supplied
gypsum silt and sand were inundated, effectively shutting
off the supply of gypsum for transport. Conversely, a moderate loss of water volume and a lowering of lake level
resulted in large muddy areas exposed to deflation on therelatively flat floor of the basin, producing a source of gypsum silt and sand. Thus, variation in gypsum flux during
overall highstands of the lake are believed to representchanges in lake volume and area.
Better age control is needed, but the evidence so far shows169
lowstands extensive areas of mudflats were available forthe production and transport of gypsum silt and sand. A
reasonable interpretation for the millimeter-scaleinterannual to decadal changes in gypsum flux therefore isthat they represent changes in gypsum transport, which
ultimately reflects wind intensity, rather than changes inlake area and volume. If this hypothesis is corroborated byadditional research, Estancia sediments may provide a rare
record of changes in paleo-wind intensity over portions of
the last deglaciation.
Geomorphic expression of climatic change
Late Pleistocene shorelinesthat cycles in gypsum flux of 100-400 years were recorded
shortly after the lake rose to its maximum extent during thelast glacial maximum. Later during the glacial-maximum
highstand, only longer cycles of 500-600 years were recorded. Oscillations in gypsum flux of less than 200 years
are preserved in the highstand sequence near 13,000 yrs
B.P.
Evidence for decadal-scale climatic changes
Bioturbating organisms were eliminated from bottom
muds near the center of the basin during prolonged
lowstands of the lake. As a result, the processes that carried
and deposited gypsum silt and sand in the lake left theirsignature as millimeter-scale structures in the lacustrine
sequence. The laminated facies is characterized by discrete,nearly pure layers of gypsum silt and sand which areseparated by laminae of carbonate-enriched clay in theupper part of the exposure. If one assumes a fairly constantsedimentation rate during accumulation of the laminatedfacies, the thickness of successive gypsum laminae provides
an approximate measure of changes in gypsum flux at thesite of deposition. A smoothed plot of consecutive gypsum
laminae thicknesses (Fig. 3B) shows the variability in
gypsum accumulation at a particular locality (playa in secs.
29 and 30, T5N, RI0E) after the recession of the glacialmaximum lake. A power spectrum of the time-series
identifies periods of 47 and 82 laminae.
Interpolation of accumulation rates between radiocarbon
dates indicates that the periods of the largest spectral peaks
in Fig. 3B correspond to cycles of —50 and —90 years in
gypsum flux. This estimate is approximate, but serves toshow that variations in the flux of gypsum occurred at alltime scales, whereas only lower-frequency changes arepreserved in highstand intervals due to disruption of therecord by bioturbation. In contrast to highstands, during
prolonged

Meinzer (1911) described the occurrence of numerous
wave-cut cliffs, beaches, beach ridges, spits, and bars associated with ancient Lake Estancia. The continuous recordof climatic and hydrologic changes is thus supplemented bygeomorphic features which provide information about theabsolute depth of the lake at various times.
All of the shoreline features described by Meinzer (1911)
occur well below the level of the topographic sill (1932 m),suggesting that the lake remained closed throughout its history. Subsequent workers, including Titus (1969), haveidentified geomorphic features near the elevation of the sill,
which are cited as evidence for a lake that overflowed thebasin. If these features are lacustrine in origin, then they mustbe much older than the last glacial-maximum highstand of
Lake Estancia (Hawley et al., 1982), as evidenced by the
degree of pedogenic-carbonate development in accompanying soils. Bachhuber (1989) suggested that the shorelines
in question may correspond to a lake that overflowed the sill
during the Illinoian glacial episode, while Hawley (pers.
comm. 1990) finds no evidence of overflow in the sill area
(Fig. I) during the Quaternary.
An absolute chronology for the definite Estancia shorelines has not as yet been established, although a reasonable
correlation between shoreline groups and the late Pleistocene
lacustrine sequence has been proposed by Bachhuber (1989)
based upon facies relationships in exposures of littoral lake
sediments. The highest shorelines correspond to the glacialmaximum highstand of Lake Estancia and consist of 12 to 14
recessional beaches and associated beach bars (Harbour,1958). The highest definite shoreline in this group occurs at
an elevation of about 1890 m, encompassing an area of 1120
km2 ("high shoreline" in Fig. 1). The highest shorelines are
fairly well preserved around the entire margin of the lakebasin and are associated with the variety of littoral landforms,
such as wave-cut cliffs and spits, described by Meinzer(1911). Good exposures of the beach sediments comprising
this shoreline group can be found in numerous gravel pits on
the western side of the basin.
The highstand of Lake Estancia that occurred near 13,000
yrs B . P. is thought to have reached an elevation of about
1870 m, covering an area of 610 km2 (Bachhuber, 1989).
Shorelines at this elevation are poorly preserved, possibly
because they would have developed on fine-grained sediments deposited during the previous highstand rather than on
the alluvial gravels of the higher shorelines. Linear shorelinefeatures are visible on air photos in the southeastern part ofthe basin at about 1871 m, where a large gravel spit from theglacial-maximum lake extends into the basin.
The lowest distinct shoreline feature in the Estancia Valley
consists of a linear beach bar occurring at about 1860 m

170(Fig. 1), or 8 m above the valley floor. This feature iscomposed of gypsiferous material that was clearly derived
from numerous small gypsum dunes that postdate the perennial lake. Thus, the gypsum beach bar indicates that ashallow lake reoccupied the basin subsequent to the initial
desiccation of Lake Estancia and the formation of the gypsum-dune field.
Ground-water fluctuation and wind activity:
the Holocene record
Desiccation of Lake Estancia between 12,000 to 11,000
yrs B.P. is preserved stratigraphically as —50 cm of homogeneous gypsiferous sediment that overlies and fills in
mudcracks that extend into laminated lake sediments at the
top of the lacustrine sequence. This white, partially indurated stratum comprises the central floor of the Estancia
Valley and marks the level of the water table upon desiccation of the lake.
Two compositionally and genetically distinct dune features rest on the Estancia Valley floor. The first set todevelop consists of relatively small dunes (<4 m high)composed of gypsiferous sediment apparently derived fromgypsum silt and sand accumulations associated with thedesiccation of Lake Estancia. The gypsum dunes werereworked by wave and current action to form the gypsum
beach bar at 1860 m. No gypsum-sand dunes occur outside
the area circumscribed by the beach bar, suggesting thatdune migration may have been halted by the rise in the lake
and that subaqueous conditions not only modified theinterior dunes into their present, hummocky form, but alsostabilized them against further migration.
The second, and most spectacular, dune structures on thevalley floor developed during a period of extensive winddeflation of previously deposited lake sediments. The material removed by deflation accumulated along the leeward
side of the deflation basins to form a series of large dunes(lunettes) that rise up to 40 m above the valley floor. Theselarge dunes are composed primarily of reworked late Pleistocene lake sediments and contain a much higher
proportion of clay than the older gypsum dunes.
Bowler (1973) studied the processes responsible for claypelletization, deflation, and the formation of clay dunes,
and among other factors cites the necessity of strong seasonality in temperature, a saline environment, and the presence of water at or near the surface. The formation anddeflation of clay pellets in the Estancia playas is no longer
occurring on a large scale, although the requirements ofshallow, saline water are met, suggesting that the clay
dunes may represent past periods of stronger seasonality
and higher summer temperatures.
Cursory examination of the blowouts and their associatedclay dunes indicates that deflation may have proceeded during at least two distinct episodes, as evidenced by large-scale
truncation surfaces that cut both earlier clay dune deposits
and the late Pleistocene lake sediments. An additional featureof the Estancia blowouts, as noted by Meinzer (1911), is thecommon presence of a well-developed bench occurring a few
meters above the level of the modern playa floors. Thisfeature may represent a gradation to base level during a brieflake stand between episodes of deflation or subsequent to
deflation. Since the period of wind deflation, 1 to 2 m ofaggradation and accumulation of playa sediments has
occurred on the floors of the blowouts, signifying a net rise in
the water table in the area of the playas.

Very little is known about the absolute timing of the postglacial geomorphic evidence for climatic and hydrologicchange that exists in Estancia Valley. Inferred changes in
the water table for the past 20 kyrs are depicted diagrammatically in Fig. 4, which assumes that: (I) the brief lake
stand represented by the gypsum beach bar occurredshortly after desiccation of Lake Estancia, and (2) deflation
occurred primarily in the mid-Holocene. The timing andmagnitude of the inferred ground-water fluctuations, andthe shape of the curve in Fig. 4, will undoubtedly change
somewhat as more dates and further geological information
become available.
Acknowledgments
Support for my research in Estancia Valley has been provided by the National Science Foundation (grant EAR8903840), the New Mexico Geological Society, and Sigma
Xi. I am grateful to L. V. Benson and J. P. Bradbury for
their reviews of this manuscript, and to R. Y. Anderson andJ. W. Hawley for sharing their thoughts on the Estancia
record.
References
Allen, B. D., and Anderson, R. Y., 1989, Decadal to millennial climate
variability in lacustrine deposits from central New Mexico (abs.): EOS,
v. 70, p. 1129.
Antevs, E., 1935, Age of the Clovis lake clays: Philadelphia Academy of
Natural Sciences, Proceedings, v. 87, pp. 304-312.
Antevs, E., 1954, Climate of New Mexico during the last glacio-pluvial:
Journal of Geology, v. 62, pp. 182-191.
Bachhuber, F. W., 1971, Paleolimnology of Lake Estancia and the Quaternary history of the Estancia Valley, central New Mexico: Unpublished
PhD dissertation, University of New Mexico, Albuquerque, 238 pp.
Bachhuber, F. W., 1989, The occurrence and paleolimnologic significanceof cutthroat trout (Oncorhynchus clarki) in pluvial lakes of the Estancia
Valley, central New Mexico: Geological Society of America, Bulletin, v.
101, pp. 1543-1551.
Bowler, J. M., 1973, Clay dunes: their occurrence, formation and environmental significance: Earth Science Reviews, v. 9, pp. 315-338.
Cody, R. D., and Cody, A. M., 1988, Gypsum nucleation and crystal
morphology in analog saline terrestrial environments: Journal of Sedimentary Petrology, v. 58, pp. 247-255.
Darton, N. H., 1928, "Red beds" and associated formations in New Mexico:
U.S. Geological Survey, Bulletin 794, 356 pp.

171
DeBrine, B. E., 1971, Quantitative hydrologic study of a closed basinwith a playa (Estancia Valley, New Mexico): Unpublished PhD
dissertation, New Mexico Institute of Mining & Technology, Socorro,
165 pp.
Harbour, J., 1958, Microstratigraphic and sedimentation] studies of an
Early Man site near Lucy, New Mexico: Unpublished MS thesis, University of New Mexico, Albuquerque, 111 pp.
Hawley, J. W., 1986, Physiographic provinces; in Williams, J. L. (ed.),
New Mexico in maps: University of New Mexico Press, Albuquerque,
pp. 23-27.
Hawley, J. W., Foster, R. W., Broadhead, R., and Love, D. W., 1982,
Road-log segment 1-B: Tijeras Canyon to Abo Canyon via Estancia andManzano: New Mexico Geological Society, Guidebook 33, pp. 8-24.
Jenkins, D. N., 1982, Geohydrology of the Madera Group, western Estancia basin, New Mexico: New Mexico Geological Society,
Guidebook 33, pp. 361-366.
Leopold, L. B., 1951, Pleistocene climate in New Mexico: American
Journal of Science, v. 249, pp. 152-168.
Meinzer, 0. E., 1911, Geology and groundwater resources of Estancia
Valley, New Mexico: U.S. Geological Survey, Water-Supply Paper
275, 89 pp.
New Mexico State Engineer Office, 1965, Chemical characteristics ofground water, Estancia underground water basin: New Mexico StateEngineer, Open-File Map.
New Mexico State Engineer Office, 1967, Thickness of alluvium, valley

fill and lake deposits, Estancia underground water basin: New Mexico
State Engineer, Open-File Map.
Pennak, R. W., 1989, Fresh-water invertebrates of the United States: John
Wiley and Sons, Inc., New York, 628 pp.
Smith, L. N., and Anderson, R. Y., 1982, Pleistocene—Holocene climateof the Estancia basin, central New Mexico: New Mexico GeologicalSociety, Guidebook 33, pp. 347-350.
Smith, R. E., 1957, Geology and ground-water resources of Torrance
County, New Mexico: New Mexico Bureau of Mines & Mineral Resources, Ground-Water Report 5, 186 pp.
Thompson, R. S., Toolin, L. J., Forester, R. M., and Spencer, R. J., 1990,
Accelerator-mass spectrometer (AMS) radiocarbon dating of Pleistocene lake sediments in the Great Basin: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 78, pp. 301-313.
Titus, F. B., 1969, Late Tertiary and Quaternary hydrogeology of Estancia
basin, central New Mexico: Unpublished PhD dissertation, Universityof New Mexico, Albuquerque, 179 pp.
Waters, M. R., 1989, Late Quaternary lacustrine history and paleoclimatic
significance of pluvial Lake Cochise, southeastern Arizona: Quaternary
Research, v. 32, pp. 1-11.
Wells, S. G., Anderson, R. Y., McFadden, L. D., Brown, W. J., Enzel, Y.,
and Miossec, J.-L., 1989, Late Quaternary paleohydrology of theeastern Mojave River drainage, southern California: New Mexico
Water Resources Research Institute, Report 242, 253 pp.

Summary of recent work on the geomorphology of the upper Rio Puerco
Paula J. Slavin
Department of Geology, University of New Mexico, Albuquerque, New Mexico 8713/

Introduction
The Rio Puerco is an entrenched meandering river of highsuspended load that heads in the Nacimiento Mountains ofnorth-central New Mexico and flows approximately 250 km
to its confluence with the Rio Grande near Bernardo, New
Mexico. Many investigators have concentrated on the geomorphology of the lower reaches of the Rio Puerco (Wells etal., 1982; Love and Young, 1983; Love, 1986), but little
work has been done on the geomorphology of the upper
reaches of the river.
The Rio Puerco, in its northern reaches, encounters theJemez lineament near the town of San Luis, New Mexico.
The Jemez lineament was first defined by the alignment of
volcanic features stretching from southwestern Arizonathrough New Mexico to southeastern Colorado (Mayo,
1958). Subsequent investigations have demonstrated astructural basis for the lineament, suggesting that its roots
extend to the mantle and that it may be a reactivated
Precambrian feature (Lipman and Mehnert, 1979; Ander,
1980; Lipman, 1980; Aldrich and Laughlin, 1984). The
effect of the lineament on the geomorphic development of
the Rio Puerco is presently under investigation by the author
in partial fulfillment of the requirements for a Master of
Science degree at the University of New Mexico. A briefsummary of some of this work follows.
The MS thesis study area extends from La Ventana
south to Mesa Prieta. Several geomorphic surfaces havebeen described in the area, ranging in age fromapproximately 2.2 m.y. to the present. One set of thesesurfaces is described below.
Rio Puerco volcanic necks
The Jemez lineament in New Mexico is defined by basalt
flows and volcanic necks aligned from southwestern tonortheastern New Mexico. In the vicinity of San Luis, the
20 to 30 volcanic necks are collectively termed the Rio
Puerco necks. The most famous of these is Cabezon Peak,visible for nearly 100 miles in all directions because of its
mammoth size, rising 538 m above the floor of the Rio
Puerco.
A comparison of the heights of the Rio Puerco necks
above base level (the Rio Puerco's channel) is presented inFig. I. It is assumed that the tops of the necks represent
former positions of the Rio Puerco valley floor and that the
Mesozoic bedrock (shales and sandstones) and alluvial ma-

172terial through which the basalt was extruded were subsequently stripped away by the river.
It appears possible that four surfaces (i.e. former valley
floors) were developed over time by the Rio Puerco as cut
down through its alluvium and Mesozoic bedrock. The
highest surface, S I , ranging in height from 493 to 566 m(1618-1857 ft) above base level, and defined by Mesa
Prieta, Cabezon Peak, and Cerro Jacobo, may beapproximately 2.2 m.y. old (Armstrong et al., 1976). The
next surface, S 2 , ranges in height from 391 to 420 m (12811378 ft) above base level. The most striking surface, S3, atapproximately 300 m (985 ft) is defined by seven of the 17
Rio Puerco necks examined (e.g. Cerro de Guadalupe,
Cerro Tinaja, Cerrito Cochin). It ranges in elevation from
297 to 323 m (974-1060 ft) above base level. The fourth
surface, S4, is defined on the basis of two necks, and isapproximately 200 m (655 ft) above base level. One of the
S4 necks (Cerro Chafo) has been sampled for dating byFrank Perry of the University of New Mexico (results
pending). If a date for this neck is obtained, an average
incision rate can be calculated for the Rio Puerco from thetime of the highest surface to the lowest surface.
Implications for the geomorphology of the Rio Puerco
The Rio Puerco is presently in a phase of incising in its
northern reaches. Photographic evidence from the late 1880's
indicates that the river has incised up to 15 m (50 ft) in 100
years, an extremely rapid incision rate. Paleochannels within
the walls of the river's channel indicate that it has undergone
many cycles of cutting and filling. It is possible that theincision of the Mesozoic shales and alluvial material surrounding the Puerco necks was similarly rapid and that an
extraordinarily fast base-level change is recorded in the surfaces of the Rio Puerco valley. It is conceivable that 400 m or
more (>1312 ft) of bedrock and alluvial material were
removed over a very short span of time (i.e. tens of thousands
to only thousands of years, based on the above incision rate
for the alluvial material eroded over the past 100 years). Even
if the erosion of 400 to 500 m (1312-1640 ft) of material has
occurred over as much as the last 2.2 m.y.,these postulated incision rates may force a reconsiderationof previously held notions that geologic processes occur
slowly over long periods of time.
Acknowledgments
I would like to thank the following people for their supportand interest in this project: S. G. Wells, F. V. Perry, C. D.
Harrington, L. D. McFadden, L. A. Woodward, D. Love,
and J. Hawley. This work has benefitted greatly by financial
and material support from the Los Alamos National Laboratory, the New Mexico Geological Society, the Bureau ofLand Management, and the University of New Mexico Student Research Allocations Committee.
References
Aldrich, M. J., Jr., and Laughlin, A. W., 1984, A model for the tectonic
development of the southeastern Colorado Plateau boundary: Journal ofGeophysical Research, v. 89, no. 10, pp. 10,207-10,218.
Ander, M. E., 1980, Geophysical study of the crust and upper mantle
beneath the central Rio Grande rift and adjacent Great Plains and Colorado Plateau: Unpublished PhD dissertation, University of New
Mexico, 218 pp.
Armstrong, R. L., Speed, R. C., Graustein, W. D., and Young, R. L.,
1976, K—Ar dates from Arizona, New Mexico, Montana, Nevada,Utah, and Wyoming: Isochron/West, no. 16, pp. 1-16.
Lipman, P. W., 1980, Cenozoic volcanism in the western United States:
Implications for continental tectonics; in Continental tectonics: National
Academy of Sciences, Washington, D.C., pp. 161-174.
Lipman, P. W., and Mehnert, H. H., 1979, Potassium—argon ages from
the Mount Taylor volcanic field, New Mexico: U.S. Geological Survey,
Professional Paper 1124-B, pp. 131-138.
Love, D. W., 1986, A geological perspective on sediment storage anddelivery along the Rio Puerco, central New Mexico; in Hadley, R. F.
(ed.), Drainage basin sediment delivery: International Association of
Hydrological Sciences, Publication 159, pp. 305-322.
Love, D. W., and Young, J. D., 1983, Progress report on the late Cenozoic
evolution of the lower Rio Puerco: New Mexico Geological Society,
Guidebook 34, pp. 277-284.
Mayo, E. B., 1958, Lineament tectonics and some ore districts of theSouthwest: American Institute of Mining Engineers, Transactions, v.
211, pp. 1169-1175.
Wells, S. G., Bullard, T. F., Condit, C. D., Jercinovic, J., Lozinsky, R. P.,
and Laughlin, D. E., 1982, Geomorphic processes on the alluvial valley
floor of the Rio Puerco: New Mexico Geological Society, Guidebook
33, pp. 45-47.

New Mexico Bureau of Mines & Mineral Resources Bulletin 137, April 1991

173

Lower and Middle Pennsylvanian stratigraphic relations, type Derryan
region, southern New Mexico and westernmost Texas
William W. Clopine¹, Walter L. Manger2, Patrick K. Sutherland', and David A. Kaiser''Conoco Inc., P.O. Box 51266, Lafayette, Louisiana 70505; 'Department of Geology, University of Arkansas, Fayetteville, Arkansas 72701;'School of Geology and Geophysics, University of Oklahoma, Norman, Oklahoma 73019; 4Anadarko Petroleum, 211 N. Robinson, Oklahoma City, Oklahoma 73102

Introduction
Thompson (1942) proposed the name Derry as a series
to include all rocks between the basal Pennsylvanian
System and the Desmoinesian Series in central and southcentral New Mexico. At its type section in the Derry Hills(Stop 1), strata of the Derryan Series rest unconformablyon the Upper Devonian Percha Formation, and Thompson(1942) stated that nowhere in the region was the base of theseries conformable. Precise biostratigraphic limits were notassigned to the series, although Thompson (1942, p. 29)
stated that unspecified species of the fusulinids
Profusulinella and Eoschubertella could be used torecognize and zone the Derryan below the Zone of
Fusulinella. He reported Millerella and Ozawainella(?)
from the beds below the first appearance of Profusulinella,but he did not discuss the relationship of the base of the
Derryan Series to the Morrowan Series. That omission wasaddressed by assigning the Derryan a post-Morrowan andpre-Desmoinesian age, as part of the Middle Pennsylvanian(Thompson, 1948, p. 69), but the series received little
attention outside of New Mexico.
Spivey and Roberts (1946, p. 186) rejected the nameDerryan in favor of their new Atokan Series because thetype Derryan section was much thinner than typical Atokan
strata, although no type section for the Atokan Series was
proposed. Moore and Thompson (1949) utilized the names
Derryan and Atokan interchangeably to designate thelowest stage of their proposed Oklan Series, but the latter
chronostratigraphic unit has been ignored. Subsequently,
the Derryan Series has become equated with the Zones ofProfusulinella and Fusulinella and as a synonym of the
Atokan (Douglass, 1977). In spite 'of the failure of the
Denyan Series to gain widespread acceptance, its typeregion in southern New Mexico and westernmost Texas
remains a significant reference for Atokan biostratigraphy,
particularly based on fusulinids.
This field trip will examine the stratigraphic relations ofMorrowan and Atokan strata on the Robledo shelf and Orogrande basin in the type "Derryan" region, with an emphasis
on series boundaries. Papers by Clopine, Sutherland, andKaiser and Manger are included as a supplement to the stop
descriptions, to provide additional faunal information. Detailed descriptions of the measured sections comprising thethree stops visited on this field trip are available as New
Mexico Bureau of Mines & Mineral Resources Open-file
Report 373 (Clopine, 1991).
Acknowledgments
The field and laboratory studies upon which this field trip
is based were undertaken with the support of National Science Foundation grants EAR-8520684 to Manger and EAR85171591 to Sutherland. The support of the New Mexico
Bureau of Mines & Mineral Resources and the encourage

ment of Frank Kottlowski, Director, are gratefully
acknowledged. We also thank Jiri Zidek, New Mexico
Bureau of Mines & Mineral Resources, for his help inediting this manuscript and, most of all, for his patience.
Geologic setting
The Early and Middle Pennsylvanian strata under consideration reflect deposition on the broad Robledo carbonate
shelf bounded by the Pedernal uplift (east and northeast),
Zuni—Defiance arch (northwest), Florida axis (southwest),
and Orogrande basin (southeast) (Fig. 1). The axis of the
Orogrande basin had a northeast—southwest orientation during Morrowan time (Fig. 2), changing through time to its
present north—south orientation. The type Derryan section(Stop 1) lies well onto the Robledo shelf, while the Bishop
Cap section (Stop 2) is near the shelf margin and the Vinton
Canyon section (Stop 3) lies along the Orogrande basin axis.
Pennsylvanian deposition followed a regional regression
and sustained period of erosion. The Helms Formation(Chesterian), dominated by terrigenous clastics, is the
youngest pre-Pennsylvanian unit involved with the unconformity and its present distribution is confined to the Orogrande basin. Post-Helms carbonate deposition began withearly Morrowan phylloid algal banks and associated mixed
skeletal wackestones in the slowly subsiding Orogrande basin. The hiatus at the base of the Pennsylvanian Systemincreases northward, reflecting transgression from that basin.
Northward and westward transgression continued through theMorrowan, resulting in the onlap of shallow, subtidal

carbonate-shelf deposits onto progressively older strata. Bylate Morrowan time, transgression reached the central Robledo shelf, producing the thin algal mudstones and shales at
the type Derryan section (Stop 1). Coeval coral and/or
Chaetetes boundstones developed in the more basinward
settings seen at Bishop Cap (Stop 2) and Vinton Canyon(Stop 3). Terrigenous clastic influx was confined to thinquartz sands and silty shales derived from the Pedernal
uplift, scattered unpredictably through the carbonate sections.
There is no apparent depositional break at the Morrowan/
Atokan boundary within the Orogrande basin, but an apparently minor unconformity occupies that position on thebasin margin and Robledo shelf (type Derryan section, Stop
1; Bishop Cap section, Stop 2). Continued transgression
during the lower Atokan finally covered the Robledo shelf,
but relative subsidence rates remained higher in the Orogrande basin, where carbonate deposition kept pace withsubsidence. Thus, the lower Atokan Series is represented bya blanket-like sequence of shallow subtidal, intertidal, andsupratidal carbonates across the entire shelf. During middle
Atokan time, both the Zuni—Defiance arch and Pedernal
uplift began to contribute terrigenous clastics that resulted inthicker and more frequent shale intervals in the Orograndebasin and central Robledo shelf. Nevertheless, shallow-shelfcarbonate deposition continued into Desmoinesian time,with the same pattern of variable subsidence rates reflectedby thick sequences in the Orogrande basin and its margin,
and thinner sequences on the Robledo shelf.
Lithostratigraphy
Pennsylvanian lithostratigraphy has seen the proposal ofa plethora of names, yet little clarification of its relationships; a situation that remains unresolved. Fig. 3 illustrates

the various lithostratigraphic classifications proposed byprevious workers in the region.
Gordon (1907) proposed the name Magdalena Group for
Pennsylvanian strata exposed in the Magdalena mining district of Socorro County, New Mexico. As defined, it encompassed the entire sedimentary succession in central New
Mexico between the Mississippian and Permian Systems.
Gordon (1907) incorporated names already used by Herrick(1900) and Keyes (1903) as subdivisions of the Magdalena
Group: Sandia Formation for a basal terrigenous clastic
unit, and the succeeding Madera Limestone. Needham(1937) rejected the usage of Sandia and Madera as
formations on the grounds that they could not be recognizedconsistently, thus reducing the Magdalena to formational
status. He also provided the first account of fusulinids fromthese rocks in New Mexico.
The terrigenous clastic lithologies developed in the type
Magdalena region are absent in the more completely carbonate sequences in southern New Mexico and westernmost
Texas. That relationship led Nelson (1940) to subdivide the
Magdalena Formation in the northern Franklin Mountains,
west Texas, informally into the La Tuna, Berino, and Bishop
Cap members (Fig. 3). While no criteria for differentiating
those members were provided, later workers, including
Thompson (1948), Harbour (1958), Lane et al. (1972), Seager(1973), and Lane (1974), used Nelson's names as either
members or formations.
Thompson (1942) rejected all previous lithostratigraphic
classifications for the Pennsylvanian strata of central andsouthern New Mexico because they lacked adequately described boundaries. He proposed a new nomenclature involving one new series, eight new groups, fifteen newformations, and one new member, based mainly on measured sections in the Mud Springs Mountains and Derry
Hills, Sierra County, New Mexico (Fig. 3). Type sections
for many of Thompson's units were poorly defined and unit
boundaries were not clearly differentiated. While Thompson's nomenclature received some limited acceptance(Gehrig, 1958; Kottlowski, 1960; King, 1973), none of thesubdivisions is currently utilized by the New Mexico
Bureau of Mines & Mineral Resources.
Kelley and Silver (1952) argued that Thompson's lithostratigraphic divisions were actually faunal zones and not
mappable units. They proposed three new names for Pennsylvanian rocks in the Caballo Mountains, Sierra County,
New Mexico, to replace those of Thompson. These are, in
ascending order: Red House, Nakaye, and Bar B (Fig. 3).
No faunal evidence was used in their definition, formation
contacts were poorly defined, and these units have gained
little acceptance.
Biostratigraphic studies reported in this volume are basedon detailed measured sections (Clopine, 1991), but they
have not addressed the problems with lithostratigraphic nomenclature for the Pennsylvanian System. Consequently,
formal lithostratigraphic divisions are not employed and all
correlations suggested herein are based on biostratigraphic
data.
Biostratigraphy
Fusulinids provide the basic biostratigraphic framework
for analysis of these Morrowan—Desmoinesian sections insouthern New Mexico and westernmost Texas. Clopine(this volume) describes the general biostratigraphic succession, which includes the Zones of Millerella (Morrowan),Eoschubertella, Profusulinella, Fusulinella (Atokan), andFusulina (=Beedeina) (Desmoinesian). The Zone of Profusulinella may be further subdivided into a lower interval
characterized by P. copiosa, a middle interval with P. sp., andan upper interval containing P. munda. The Zone of
Fusulinella may be further subdividcd into a lower F. acuminata and upper F. devexa intervals. Thc fusulinid succession in southern New Mexico and westernmost Texas closely
approximatcs that of the North American midcontinent.
The basal 3.7 ft of the type Derryan section is of Morrowan age based on brachiopods reported by Suthcrland (thisvolume). That conclusion does not conflict with thc associated fusulinids. Reports of Profusulinella near the base of
the type Derryan section (Lane et al., 1972) are reinterpreted
to be Eoschubertella. The Morrowan/Atokan boundary attype Derryan is an unconformity, which also occurs in the
Bishop Cap section. In both sections, thc conodont Diplognathodus coloradoensis, indicativc of upper lower Atokan
strata, either occurs with Eoschubertella or succecds its
appearance within a fcw feet.
Fusulinids from the remaining Derryan succession correlate with the midcontinent and indicate equivalence to the
Atokan Series. The upper boundary of the type Derryan
section falls just below the first occurrencc of Fusulina,
which corresponds to the generally accepted Atokan/Desmoinesian boundary. Conodonts from thc Dcrryan section
above the appearance of Diplognathodus are dominated by
Idiognathodus and do not reflect the Atokan/Desmoinesianboundary.
Regional correlations
Figs. 4-6 illustrate correlation of fusulinid zones between
the sections examined as stops on this field trip. Fig. 4

documents thc northwestward transgression of Morrowanseas from the Orogrande basin and the onlap ofprogressively younger Morrowan beds on strata of thc
Robledo shelf, although some consideration must be given
to Morrowan/ Atokan boundary erosion to explain thinningof Morrowan strata. Figs. 5 and 6 reflect the continued
transgression of Atokan and lower Desmoinesian seas thatdeveloped the blanket-like character of these strata in theregion. Differential subsidence of the Orogrande basincompared to the Robledo shelf is illustrated by comparative
thicknesses of the type Derryan section and those at Bishop
Cap and Vinton Canyon.
Stop descriptions
Stop 1: Type Derryan section, Sierra County, New
MexicoLocation: This section is a roadcut along 1-25 located in
the Derry Hills, just north of the Dona Ana County line,
approximately 0.8 mi east of Derry, New Mexico (sec. 29,
T175, R4W, Garfield 7'/2' quadranglc). This exposure includes Thompson's (1942) section 19.
Significance: Thompson (1942) proposed the scction
exposed in the draw on the northwest end of this roadcut as
the type section for the Derryan Series. The lithologic
succession can be easily matched to his dcscription, although the formation names Thompson proposed are notrecognizable. Fig. 7 illustrates the general lithologic
succession and fusulinid occurrences at this section.
Of particular interest here is the excellent exposure of thebasal Pennsylvanian boundary relations. "Derryan" strata reston sandstone and shale of the Upper Devonian Percha
Formation. Basal Morrowan strata comprise approximately

2 ft of shaly, nodular wackestone grading upward into 1 ft of
nodular wackestone. A typical upper Morrowan brachiopod
assemblage, including Zia novamexicana, Spirifer goreii,and Neochonetes platynotus, has been recovered from thisinterval together with nondiagnostic but presumably Morrowan foraminifers (see Clopine, this volume, and Sutherland, this volume).
Above the wackestone, 0.7 ft of gray to tan clay, with finequartz silt and plant(?) fragments, may represent a paleosol
developed at the Morrowan/Atokan boundary. Conodontsfrom this unit are obviously reworked and include
Idiognathodus and Adetognathus. The light clay is overlain
by 0.8 ft of fissile black shale that contains horizontal gypsum
laminae and lenses. Cleavage is well developed in the lower
clay, but the cleavage planes do not pass through thesucceeding black shale. In contrast, the black shale is highly
fractured, but the fractures do not correspond to structurescontinuing from the underlying shale (Fig. 8).
Carbonates succeeding the black shale contain Diplogna

thodus and Eoschubertella(?). Profusulinella appears at thebase of unit 23, at a horizon 44 ft above the appearance of
Eoschubertella(?) and 44.8 ft above the Morrowan/Atokan
unconformity. That occurrence is the only Profusulinella
from the type Derryan. Fusulinella appears in unit 25, 85 ft
above the base of the section. The Atokan/Desmoinesianboundary, marked by Fusulina, appears in unit 60, 125 ft
above the base of the section.
The entire type Derryan section is dominated by the Idiognathodus biofacies (see Kaiser and Manger, this volume).
That taxon was the most frequently occurring element, followed by Neognathodus. Recoveries varied from barren, ordepauperate to lag concentrations of hundreds of ramiforms,
particularly from units 15, 19, 38, and 43.
Stop 2: Bishop Cap section, Dona Ana County,
New Mexico
Location: This section is located on the steep north-facing
slope of Pyramid Peak, 1 mi north of Bishop Cap, approx-

imately 15 mi southeast of Las Cruces, New Mexico (sec. 23,
T24S, R3E, Bishop Cap 71/2' quadrangle). New Mexico
Bureau of Mines & Mineral Resources Geologic Map 29(Seager, 1973) provides a composite stratigraphic column,
geologic history, and geologic map of the area.
Significance: The Bishop Cap section occupied a shelfmargin position, and its thicker sequence of carbonate lithologies tends to be more grain-dominated and fossiliferous
than those seen at the type Derryan section (Figs. 9, 10).
Corals are particularly common in the lower portion of thissection, and of particular interest here is a boundstone of thecolonial rugose coral Petalaxis (unit 29). Sutherland (1985)
noted three other occurrences of Petalaxis similar to the one
at Bishop Cap: southwestern Ozark region, northeastern
Oklahoma; Ardmore basin, south-central Oklahoma; and the
Llano uplift, central Texas. All North American
midcontinent Petalaxis localities seem to fall within the upper Morrowan Idiognathodus sinuosus conodont zone andmay in fact represent precisely the same age, although thegenus ranges into strata equivalent to the DesmoinesianSeries elsewhere. The Petalaxis boundstone succeeds aninterval with a thin arkosic sandstone (unit 21).
The Mississippian—Pennsylvanian boundary is not particularly well exposed at this locality, but Morrowan conodonts
occur through the basal 138 ft of the Bishop Cap section inassociation with nondiagnostic foraminifers, including Millerella. Diplognathodus coloradoensis was recovered fromunit 35 and the first Eoschubertella was found in unit 38, with
Profusulinella appearing in unit 39, only 1.2 ft higher. Thus,
the interval that could possibly represent the zone of
Eoschubertella is quite thin, and the Morrowan/Atokanboundary is probably an unconformity here as well as at thetype Derryan section (Stop 1), since Diplognathodus is typicalof an upper lower Atokan position.
Fusulinella makes its appearance in unit 81, 283 ft above thebase of this section, followed by Fusulina, marking the
Atokan/Desmoinesian boundary in unit 117, 398 ft above thebase of the section. Conodont recoveries from the Mor-rowan interval are poor, and the Adetognathus biofacies
dominates that portion of the Bishop Cap section. Idiognathodus appears in samples from unit 17, with Neognathodus bassleri, but conodonts do not become common until
the level of unit 70. From that horizon through theremainder of the Atokan section, lag concentrates oframiform elements, dominated by Idiognathodus and
Neognathodus, are typical, particularly in units 93, 95, 100,
and 111. The Idiognathodus biofacies characterizes the
Atokan interval at Bishop Cap.

Stop 3: Vinton Canyon section, El Paso County, Texas
Location: Vinton Canyon is located on the west side of
the northern Franklin Mountains, approximately 23 mi southof Las Cruces, New Mexico, and 15 mi north of El Paso,
Texas (Canutillo 7'/2' quadrangle). Nelson (1940) proposed
the La Tuna, Berino, and Bishop Cap members of the Magdalena Formation from the Vinton Canyon section. Lane
(1974) included conodont data from this section in his analysis of the wedge-on-wedge relationships in the Mississippian and Lower Pennsylvanian sequences of the region.
Significance: Vinton Canyon preserves the thickest andleast variable succession of Pennsylvanian strata seen on thefield trip (Fig. 11). The section always occupied a position along the axis of the Orogrande basin. As subsidence
kept pace with deposition, Pennsylvanian carbonates are
dominated by algal biomicrites, with wackestones being thenext best represented lithology; grainstones are thin and rarely
developed.
The contact of Pennsylvanian carbonates with the Helms
Formation (Chesterian) is particularly well exposed at Vinton Canyon, but its placement is somewhat problematic
because of poor fossil recoveries through similar-appearing
lithologies. The base of the Pennsylvanian System is placedat the base of unit 6 (Fig. 12).
Morrowan algal biomicrites and wackestones with Millerella and impoverished conodont recoveries extend to theappearance of Eoschubertella in unit 58, 246 ft above the baseof the Pennsylvanian. Profusulinella appears in unit 62, 11 ft
higher. Neither fusulinid appearance is accompanied bydiagnostic
conodonts.
No
unconformity
attheMorrowan/Atokan boundary is indicated or suspected at
Vinton Canyon.
The appearance of Fusulinella occurs in unit 119, 446 ft
above thc base of the Pennsylvanian section, followed by
Fusulina at the Atokan/Desmoinesian boundary in unit 164,674 ft above the base of the section. Samples processed for
conodonts from those levels were essentially barren but
contained Idiognathodus, suggesting that the Atokan portionof the Vinton Canyon section also falls within the biofacies
characterized by that taxon.

181

References
Clopine, W. W., 1991, Lower and Middle Pennsylvanian measured sections, type Derryan region, southern New Mexico and westernmost
Texas: New Mexico Bureau of Mines & Mineral Resources, Open-file
Report 373, 121 pp.
Connolly, W. M., and Stanton, R. J., 1986, Lower Pennsylvanian (Morrowan) sedimentation in the Orogrande Basin; in Southwest Section of
A.A.P.G., Transactions and guidebook of 1986 Convention, Ruidoso,
New Mexico: New Mexico Bureau of Mines & Mineral Resources, pp.
129-142.
Douglass, R. C., 1977, The development of fusulinid biostratigraphy; in
Kauffman, E. G., and Hazel, J. E. (eds.), Concepts and methods ofbiostratigraphy: Dowden, Hutchinson, and Ross, Stroudsburg, Pennsylvania, pp. 463-481.
Gehrig, J. L., 1958, Middle Pennsylvanian brachiopods from the Mud
Springs Mountains and Derry Hills, New Mexico: New Mexico Bureauof Mines & Mineral Resources, Memoir 3, 24 pp.
Gordon, C. H., 1907, Mississippian (Lower Carboniferous) formations in
the Rio Grande Valley: American Journal of Science (4), v. 24, no. 174,
pp. 58-64.
Herrick, C. L., 1900, The geology of the White Sands of New Mexico:
Journal of Geology, v. 8, pp. 112-128.
Harbour, R. L., 1958, Pennsylvanian and Permian rocks in the northernFranklin Mountains, Texas (abs.): Geological Society of America, Bulletin, v. 69, no. 12, pt. 2, pp. 1727-1728.
Kelley, V. C., and Silver, C., 1952, Geology of the Caballo Mountains:
University of New Mexico Publications in Geology, No. 4, 286 pp.
Keyes, C. R., 1903, Geological sketches of New Mexico: Ores and Metals,
v. 12, p. 48.
King, W. E., 1973, Fusulinids Millerella and Eostaffella from the Pennsylvanian of New Mexico and Texas: New Mexico Bureau of Mines &Mineral Resources, Memoir 26, 34 pp.
Kottlowski, F. E., 1960, Summary of Pennsylvanian sections in southwestern New Mexico and southeastern Arizona: New Mexico Bureauof Mines & Mineral Resources, Bulletin 66, 187 pp.
Lane, H. R., 1974, Mississippian of southeastern New Mexico and West
Texas: a wedge-on-wedge relation: American Association of Petroleum
Geologists, Bulletin, v. 58, pp. 269-282.
Lane, H. R., Sanderson, G. A., and Verville, G. J., 1972, Uppermost
Mississippian-basal Middle Pennsylvanian conodonts and fusulinids
from several exposures in the south-central and southwestern UnitedStates: Twenty-fourth International Geological Congress, Montreal,sec. 7, pp. 549-555.
LeMone, D. V., King, W. E., and Cunningham, J. E., 1974, Pennsylvanian
System of Chloride Flat, Grant County, New Mexico: New Mexico
Bureau of Mines & Mineral Resources, Circular 131, 18 pp.
Moore, R. C., and Thompson, M. L., 1949, Main divisions of Pennsylvanian period and system: American Association of Petroleum Geologists, Bulletin, v. 33, pp. 275-301.
Needham, C. E., 1937, Some New Mexico Fusulinidae: New Mexico
Bureau of Mines & Mineral Resources, Bulletin 14, 59 pp.
Nelson, L. A., 1940, Paleozoic stratigraphy of Franklin Mountains, west
Texas: American Association of Petroleum Geologists, Bulletin, v. 24,
pp. 157-172.
Seager, W. R., 1973, Geologic map and sections of Bishop Cap-Organ
Mountains area: New Mexico Bureau of Mines & Mineral Resources,
Geologic Map 29, scale 1:24,000.
Spivey, R. C., and Roberts, T. G., 1946, Lower Pennsylvanianterminology in central Texas: American Association of Petroleum
Geologists, Bulletin, v. 30, pp. 181-186.
Sutherland, P. K., 1985, The colonial rugose coral genus Petalaxis in the
Middle Carboniferous of North America and its stratigraphic
significance (abs.): Geological Society of America, Abstracts with
Programs, v. 17, no. 7, p. 730.
Thompson, M. L., 1942, Pennsylvanian System in New Mexico: New
Mexico Bureau of Mines & Mineral Resources, Bulletin 17, 92 pp.
Thompson, M. L., 1948, Studies of American fusulinids: University of
Kansas Paleontological Contributions, Protozoa, Article 1, pp. 1-184.

Lower and Middle Pennsylvanian fusulinid biostratigraphy in south-central New Mexico andwesternmost Texas: A brief review
William W. Clopine
Introduction
Lower and Middle Pennsylvanian strata in southern New
Mexico and westernmost Texas include an abundant and,for the most part, well preserved fusulinid fauna. Previous
studies of this fauna include those by Thompson (1942a,
1948), Stewart (1970), Lane et al. (1972), King (1973), andVerville and Sanderson (1986). This paper is a brief reviewof the fusulinid biostratigraphic succession in this region as
documented by those studies and by Clopine (1990).
Location of study area
The study area is located in south-central New Mexico
and westernmost Texas (Fig. 1). Four sections have beenmeasured within the study area (Clopine, 1990, 1991) along
a roughly linear northwest-southeast trend of approximately
160 km (100 mi). Pennsylvanian strata crop out along thistrend in a series of discontinuous fault-block mountains.

The measured sections are located in the northern Franklin
Mountains, El Paso County, Texas, on the northwest slope of
Pyramid Peak at Bishop Cap in Doña Ana County, New
Mexico, and on the southwest slope of the Derry Hills and in
Whiskey Canyon in the northern Mud Springs Mountains,
Sierra County, New Mexico (Fig. 1). See Clopine (1991) for
specific directions to each of these measured sections.
Zone of Millerella
The Zone of Millerella was designated by Thompson(1945, p. 40) as the lowermost Pennsylvanian fusulinid zone.
For the purpose of this paper, the Zone of Millerella ismodified from Thompson (1945, 1948) to range from the firstoccurrence of Millerella and/or Eostaffella (base of the Pennsylvanian System in the study area) to the first occurrence of
Eoschubertella and/or Pseudostaffella (i.e. the widely
accepted Morrowan Series boundary markers). The occur-36) occur in carbonate wackestone, in the absence of moreadvanced forms, 0.5 m (1.6 ft) above the base of the Pennsylvanian System. The upper limit of the Zone of Millerella
is interpreted to fall at an unconformity marking the Morrowan/Atokan boundary.
Zone of Eoschubertellarence of Eoschubertella and/or Pseudostaffella below the
first occurrence of Profusulinella (= Mamet and Skipp,
1970, Faunal Zone 21) is informally termed the zone of
Eoschubertella.
The Zone of Millerella occurs in three of the four studied
measured sections. It reaches its maximum thickness at
Vinton Canyon where it includes 75 m (246 ft) ofMorrowan strata. Northward, at Bishop Cap, the Zone ofMillerella is reduced to 44 m (143 ft). This is a thickness
loss of over 40% in 25 km (15.5 mi). Most, if not all, of the
thickness reduction in Morrowan strata from Vinton
Canyon to Bishop Cap is due to loss of lower Morrowan
strata at the basal Pennsylvanian unconformity.
Northwest of Bishop Cap, the Zone of Millerella is greatly
reduced or absent. It includes less than 1 m (3.2 ft) at the type
Derryan section. Here the Morrowan strata rest on the
Devonian Percha Shale. Millerella and Eostaffella (Fig. 2.27—

The zone of Eoschubertella is equivalent to Mamet andSkipp's (1970, p. 338) Faunal Zone 21 (basal Atokan) as usedin the Midcontinent and western United States. The zone of
Eoschubertella is recognized by the occurrence of
Eoschubertella and/or Pseudostaffella below the first occurrence of Profusulinella. This zone is recognized at each of
the sections. The lowest Eoschubertella sp. at Vinton Canyon
occurs 75 m (246 ft) above the base of the Pennsylvanian.
Rare Eoschubertella sp. occur throughout the zone which is
3.75 m (12 ft) thick at this section. Eoschubertella(?) firstoccurs 44 m (143 ft) above the base of the Pennsylvanian at
Bishop Cap. Profusulinella copiosa occurs only 0.75 m (2.5ft) above the first Eoschubertella(?). This may only reflect therarity of the latter species at this locality.
The base of the Atokan Series at the type Derryan section
falls at a minor unconformity within a poorly fossiliferous
shale interval (between units 3A and 3B). Eoschubertella(?)
first occurs 0.25 m (0.8 ft) above this unconformity and 1.2
m (3.9 ft) above the base of the Pennsylvanian (unit 3-5).
Rare unequivocal Eoschubertella (Fig. 2.26) first occurs 1.2
m (3.9 ft) higher in the section (unit 3-7). Eochubertella
mexicana (Fig. 2.25) becomes somewhat more abundant andoccurs with Pseudoendothyra sp. (Fig. 2.23-24) slightly
higher in the type Derryan section (unit 3-8). The thicknessof the zone of Eoschubertella at the type Derryan section
has been a matter of some disagreement.
Thompson (1942a, p. 33) reported Profusulinella in the"Arrey Formation" at the type Derryan section 5 m (16.5 ft)
above the base of the Pennsylvanian System (unit 3 ofThompson, 1942a; unit 3-10 of Clopine, 1990). Thompson(1948, p. 82) altered his earlier interpretation, stating thatfurther study seemed to indicate specimens from his unit 3,
formerly referred to as Profusulinella, may be a primitive
form of Eoschubertella. Thompson's latter view, supported
by data of Clopine (1990), has apparently not been noted byother workers.
Lane et al. (1972, p. 553) reported Profusulinella 4.5 m (15ft) above the base of the type Derryan section (same as
Thompson's, 1942a, lowest "Profusulinella" horizon).
However, these specimens have been re-examined by George
Verville (pers. comm. 1989), who states that they should

184
be assigned to Eoschubertella. King (1973, p. 2) reported
Profusulinella in bed 2 of Thompson (1942a; unit 3-7 ofthis author). In addition, King (1973, p. 2) suggested from"unidentifiable fragments" that the lowest Profusulinella
occurrence may be in Thompson's bed 1. He now believes,
however, that those specimens were probably Eoschubertella (William King, pers. comm. 1989).
The only documented occurrence of Profusulinella at thetype Derryan section is reported by Clopine (1990, base ofunit 3-23; unit 10 of Thompson, 1942a) 13.4 m (44 ft)
above the base of Eoschubertella(?) and 13.65 m (44.8 ft)
above the lowest Atokan strata. This interval includes all ofThompson's (1942) "Arrey Formation" and the lower 2.7 m(9 ft) of his "Apodaca Formation."
Zone of Profusulinella
The Zone of Profusulinella was designated by Thompson(1945, pp. 40-41) as the lower and middle "Derryan" (Atokan) fusulinid zone. It is bounded by the first occurrences ofProfusulinella and Fusulinella and falls between the zonesof Eoschubertella and Fusulinella. The Zone of Profusulinella has gained widespread acceptance as the lower—
middle Atokan fusulinid biozone.
The Zone of Profusulinella includes three distinct intervals in the study area, a lower interval characterized by
Profusulinella copiosa (Fig. 2.13-18), a middle interval
characterized by Profusulinella sp. (Fig. 2.8-12), and an
upper interval characterized by Profusulinella munda (Fig.
2.5-7), Profusulinella decors (Fig. 2.2-4), and Profusulinella apodacensis (Fig. 2.1a, b). The Zone of Profusulinella is 61.5 m (202 ft) thick at Vinton Canyon. Rare
Profusulinella cf. fittsi (Fig. 2.19-22) occur with P. copiosa
at the base of the zone. The zone is reduced to 41.6 m(136.5 ft) at Bishop Cap, where P. copiosa is also
associated with P. cf. fittsi. The Zone of Profusulinella isfurther reduced to approximately 8.5 m (27.9 ft) at the type
Derryan section. P. copiosa has not been recovered here orat the Whiskey Canyon section. A single horizon yielding
rare P. cf. fittsi (base of unit 3-23), however, occurs at thetype Derryan section. This horizon is tentatively correlatedwith the lower portion of the P. copiosa range zone. No
other Profusulinella have been recovered from the type
Derryan section.
The Zone of Profusulinella is recognized at Whiskey Canyon, but it is restricted to a single horizon. Thompson (1948,
p. 73) correlated this horizon (Thompson's unit 11-5; unit 4-5
of Clopine, 1990) with the upper "Apodaca Formation" at thetype Derryan section, based on his reported occurrence ofProfusulinella apodacensis. However, all fusulinids recovered from the upper "Apodaca Formation" at the type
Derryan section by Clopine (1990) have proven to be Fu

sulinella, and Thompson's correlation thus does not appear
justified.
Zone of Fusulinella
The Zone of Fusulinella was designated by Thompson(1945, p. 41) to include upper Derryan (Atokan) and lowermost Desmoinesian strata. Thompson (1948, p. 23) modified his earlier definition of the Zone of Fusulinella,lowering its upper datum to "just below the lowest
occurrence of Fusulina (Beedeina)." The base of the zone
remained unchanged at the first occurrence of Fusulinella.
The zone of Fusulinella has gained widespread acceptance as
the upper Atokan fusulinid biozone.
The Zone of Fusulinella in the study area includes alower interval characterized by Fusulinella acuminata (Fig.
3.1215) and an upper interval characterized by Fusulinella
devexa (Fig. 3.10-11). The Zone of Fusulinella reaches its
maximum thickness within the study area at Vinton
Canyon, where it is 42.18 m (138.4 ft) thick. It is reduced to9.25 m (30.4 ft) at the Bishop Cap section locality, where
Fusulinella acuminata interval does not occur due to thelack of fusulinids in the dolomitic lithologies below
Fusulinella devexa.
The Zone of Fusulinella is 15.05 m (49.4 ft) thick at thetype Derryan section. Here Fusulinella fugax (Fig. 3.1620)
occurs with Fusulinella acuminata in the upper portion ofThompson's (1942a) "Apodaca Formation." Thompson(1948) indicated the occurrence of Profusulinella apodacensis in this horizon. However, only members of the genus
Fusulinella were recovered from the upper "Apodaca Formation" by Clopine (1990). The upper boundary of the Zoneof Fusulinella corresponds to the top of the type Derryan
Series, which lies immediately below the first occurrence ofBeedeina insolita (Fig. 3.4-6). Somewhat more advanced
Beedeina (Fig. 3.2-3) and the first Wedekindellina (Fig. 3.1)
occur slightly higher in the section.
Acknowledgments
This paper is part of a University of Oklahoma PhD
dissertation supported by National Science Foundation
Grant EAR-8517591 through Patrick K. Sutherland, by agrant from Amoco Research Laboratory in Tulsa,
Oklahoma, by the University of Oklahoma School of
Geology and Geophysics, and by the New Mexico Bureauof Mines & Mineral Resources. Interpretations benefitted
from observations made in the field by Patrick K.
Sutherland, Walter L. Manger, and George J. Verville. I
would also like to thank and express my appreciation to
Patrick K. Sutherland, George J. Verville, George A.
Sanderson, Walter L. Manger, and John R. Groves for theirsignificant contributions to this study.

186
References
Clopine, W. W., 1990, Middle Pennsylvanian fusulinid biostratigraphy of
the type Derryan area in southern New Mexico and westernmost Texas:
Unpublished PhD dissertation, University of Oklahoma, Norman, 271
pp.
Clopine, W. W., 1991, Lower and Middle Pennsylvanian measured sections, type Derry region, southern New Mexico and westernmost Texas:
New Mexico Bureau of Mines & Mineral Resources, Open-file Report
373, 122 pp.
Groves, J. R., 1986, Foraminiferal characterization of the MorrowanAtokan (lower Middle Pennsylvanian) boundary: Geological Society ofAmerica, Bulletin, v. 97, pp. 346-353.
King, W. E., 1973, Fusulinids Millerella and Eostaffella from the Pennsylvanian of New Mexico and Texas: New Mexico Bureau of Mines &Mineral Resources, Memoir 26, 34 pp.
Kottlowski, F. E., 1960, Summary of Pennsylvanian sections in southwestern New Mexico and southeastern Arizona: New Mexico Bureau of
Mines & Mineral Resources, Bulletin 66, 187 pp.
Lane, H. R., 1974, Mississippian of southeastern New Mexico and West
Texas: A wedge-on-wedge relation: American Association of Petroleum
Geologists, Bulletin, v. 58, no. 2, pp. 269-282.
Lane, H. R., Sanderson, G. A., and Verville, G. J., 1972, Uppermost
Mississippian-basal Middle Pennsylvanian conodonts and fusulinids
from several exposures in the south-central and southwestern UnitedStates: Twenty-fourth International Geological Congress, Montreal,
Section 7, pp. 549-555.
Mamet, B. L., 1975, Carboniferous foraminifera and algae of the Amsden
Formation (Mississippian and Pennsylvanian) of Wyoming: U.S. Geological Survey, Professional Paper 848-B, 18 pp.
Mamet, B. L., and Armstrong, A. K., 1972, Lisburne Group, Franklin andRomanzof Mountains, northeastern Alaska: U.S. Geological Survey,
Professional Paper 800-C, pp. C127-C144.
Mamet, B. L., and Skipp, B., 1970, Preliminary foraminiferal correlationsof Early Carboniferous strata in the North American Cordillera: Congres
et Colloques, Collogue sur la Stratigraphie du Carbonifere, Universite deLiege, v. 55, pp. 325-348.
Manger, W. L., Sutherland, P. K., and Clopine, W. W., 1987, MorrowanAtokan boundary relations, type "Derryan" section, New Mexico (abs.):

Geological Society of America, South-Central Section Abstracts with
Programs, v. 19, no. 3, p. 173.
Mikhailov, A. V., 1939, K kharakteristike rodov nizhnekamennougol'nykh
foraminifer territorii SSSR: Sbomik Leningradskogo Geologicheskogo
Upravlenia, Glavnoe Geologicheskoe Upravlenie, v. 3, pp. 47-62 (notseen).
Rauzer-Chemousova, D. M., 1948, Some new species of Foraminiferafrom the Lower Carboniferous sediments of the Moscow Basin [in Russian]: Akademia Nauk SSSR, Trudy Geologicheskogo Instituta, v. 62,
no. 19, pp. 227-238.
Ross, C. A., and Sabins, F. F., 1965, Early and Middle Pennsylvanian
fusulinids from southeast Arizona: Journal of Paleontology, v. 39, pp.
173-209.
Roth, R., and Skinner, J., 1930, The fauna of the McCoy formation,
Pennsylvanian of Colorado: Journal of Paleontology, v. 4, pp. 332-352.
Stewart, W. J., 1970, Fusulinids of the Joyita Hills, Socorro County,
central New Mexico: New Mexico Bureau of Mines & Mineral
Resources, Memoir 23, part 2, pp. 32-82.
Sutherland, P. K., and Manger, W. L., 1984, Morrowan brachiopods inthe type "Derryan" Series (Pennsylvanian), New Mexico (abs.):
Geological Society of America, Abstracts with Programs, v. 16, no.
2, p. 115.
Thompson, M. L., 1935, Fusulinids from the Lower Pennsylvanian Atokan
and Boggy formations of Oklahoma: Journal of Paleontology, v. 9, pp.
291-306.
Thompson, M. L., 1936, Fusulinids from the Black Hills and adjacentareas in Wyoming: Journal of Paleontology, v. 10, pp. 95-113.
Thompson, M. L., 1942a, Pennsylvanian System in New Mexico: New
Mexico Bureau of Mines & Mineral Resources, Bulletin 17, 92 pp.
Thompson, M. L., 1942b, New genera of Pennsylvanian fusulinids: American Journal of Science, v. 240, pp. 403-420.
Thompson, M. L., 1945, Pennsylvanian rocks and fusulinids of east Utah
and northwest Colorado correlated with Kansas section: State
Geological Survey of Kansas, Bulletin 60, part 2, pp. 19-84.
Thompson, M. L., 1948, Studies of American fusulinids: University of
Kansas Paleontological Contributions, Protozoa, Article 1, 184 pp.
Verville, G. J., Sanderson, G. A., and Madsen, M. E., 1986, Pennsylvanian
fusulinids from the Fra Cristobal Range, Sierra County, New Mexico:
New Mexico Geological Society, Guidebook 37, pp. 215-223.

Morrowan brachiopods from the type "Derryan" Series (Pennsylvanian), southern New Mexico
Patrick K. Sutherland
Introduction
The Derry Series was originally defined by Thompson(1942) as including all pre-Desmoinesian Pennsylvanianstrata in southern New Mexico. Lane et al. (1972) equatedthe Derryan with the Atokan Series in Oklahoma, but theirdetailed analysis of Thompson's (1942) type section started
at the lowest cliff-forming limestone (Fig. 1, unit 5). A
subsequently developed prominent roadcut on 1-25 at thetype Derry, east of Derry, exposes the Pennsylvanian unconformity with the Upper Devonian sandstone and shale of
the Percha Formation and a 2.5 ft interval of shale andnodular limestone that is poorly exposed at the base ofThompson's (1942) type sequence.
Sutherland and Manger (1984) recovered a well-preservedlate Morrowan brachiopod fauna from the nodular limestoneand shale interval (Fig. 1, units 3a, 3b) found at the base of
the Pennsylvanian in the 1-25 roadcut, and selections fromthis fauna are here illustrated for the first time (Fig. 2). Later,
after further, detailed examination of this exposure, they
(Manger et al., 1987) observed the presence of a postMorrowan - pre-Atokan unconformity that falls within a thin
shale interval immediately above the brachiopod-

bearing units. There, a lower 0.5-0.7 ft of brown shale,possibly a paleosol (Fig. 1, unit 4a), carries reworked conodonts with upper Morrowan affinities. This unit is overlain
by 0.5-0.8 ft of barren fissile black shale (Fig. 1, unit 4b).
Thick-bedded limestones (Fig. 1, unit 5) immediately overlie this black shale and contain upper lower Atokan conodonts succeeded by Profusulinella within 13 ft (Manger etal., 1987).
Numerous thin sections of the nodular limestone (Fig. 1,
unit 3b) produced the foraminifers Millerella extensa andEostaffella sp. to the exclusion of more advanced forms(identified by J. R. Groves, pers. comm. 1985). These species are typically found in Morrowan faunas, but they also
range higher stratigraphically.
Brachiopod occurrences
The Morrowan series in the southern Midcontinent has
produced a large, varied, and highly distinctive brachiopod
fauna. Sutherland and Harlow (1973) recorded 41 Morrowan brachiopod species from the La Pasada Formation in
the southern Sangre de Cristo Mountains in northern New
Mexico. In that area, however, conodonts indicate that only

187
the uppcr part of the Morrowan Series is represented (Sutherland and Harlow, 1973). In the type Morrowan area, in
northwestern Arkansas and northeastern Oklahoma, where
the Morrowan series is more complete, Henry and Suthcrland (1977) reported the occurrcnce of about 80 brachiopod species. However, detailed descriptions andillustrations of the new clements in that fauna havc not yet
been published. Many of the species that occur in northern
New Mexico also occur in Arkansas and Oklahoma, but a
few unique species arc restricted to each area.
The 2.5 ft Morrowan interval at the base of the type
"Derryan" section has produced a late Morrowan brachiopod
fauna of 12 species that are particularly similar to theequivalcnt species that occur in thc La Pasada Formation innorthern Ncw Mexico. Of thcse, the following nine are
restricted to thc Morrowan in all known occurrences:
1. Zia novamexicana Sutherland and Harlow (Fig. 2.1a, 2.
lb); common
2. Spirifer goreii Mather (Fig. 2.3); rare
3. Neochonetes? platynotus (White) (Fig. 2.8); fairly common
4. Plicochonetes? arkansanus (Mather) (Fig. 2.9a, 2.9b);rare
5. Linoproductus nodosus (Newberry) (not figured); rare
6. Sandia welleri (Mather) (Fig. 2.4a, 2.4b); rare
7. Punctospirifer morrowensis Suthcrland and Harlow (Fig.
2.5); fairly common
8. Rhynchopora magnicosta Mather (not figured); rare
9. Composita cf. gibbosa Mather (not figured); common

188
The following three species occur in northern New Mexico in both the Morrowan and lower Atokan:
10. Antiquatonia coloradoensis (Girty) (not figured); rare
11. Anthracospirifer newberryi Sutherland and Harlow(not figured); common
12. Anthracospirifer tanoensis Sutherland and Harlow(not figured); common
The following two species, which occur in the type "Derryan" section (Fig. 1, units 3a, 3b), are not known fromother areas and possibly represent new species:
1. Rhipidomella sp. (Fig. 2.6); fairly common; larger andthicker than the Morrowan species R. trapezoids Sutherland and Harlow.
2. Punctospirifer? sp. (Fig. 2.7); rare; much larger andwith fewer plications than Punctospirifer morrowensis
Sutherland and Harlow.
Morrowan age designation
The assignment of a Morrowan age to the lower 2.5 ft
interval in the type "Derryan" sequence is a particularly
strong one. The list above includes a wide range of species
that have been found only in the Morrowan. Of particularimportance is the common occurrence of Zia
novamexicana, which forms a distinct zone in theMorrowan lower part of the La Pasada Formation in thesouthern Sangre de Cristo Mountains (see Sutherland andHarlow, 1973, fig. 7). Spirifer goreii is also a particularly
distinctive and important occurrence.

The occurrence of Morrowan fossils in the type "Derryan" sequence is important because it shows that the lower
type "Derryan" of southern New Mexico overlaps the type
Morrowan of Arkansas and cannot be regarded as theprecise equivalent of the Atokan Series.

References
Henry, T. W., and Sutherland, P. K., 1977, Brachiopod biostratigraphy ofMorrowan Series (Pennsylvanian) in northwestern Arkansas and northeastern Oklahoma; in Sutherland, P. K., and Manger, W. L. (eds.), Upper
Chesterian—Morrowan stratigraphy and the Mississippian—Pennsylvanian boundary in northeastern Oklahoma and northwestern Arkansas: Oklahoma Geological Survey, Guidebook 18: 107-116.
Lane, R. R., Sanderson, G. A., and Verville, G. J., 1972, Uppermost
Mississippian—basal Middle Pennsylvanian conodonts and fusulinids
from several exposures in the south-central and southwestern UnitedStates: Twenty-fourth International Geological Congress, Montreal,
Section 7: 549-555.
Manger, W. L., Sutherland, P. K., and Clopine, W. W., 1987, MorrowanAtokan boundary relations, type "Derryan" section, New Mexico (abs.):
Geological Society of America, Abstracts with Programs, 19: 173.
Sutherland, P. K., and Harlow, F. H., 1973, Pennsylvanian brachiopods
and biostratigraphy in southern Sangre de Cristo Moountains, New
Mexico: New Mexico Bureau of Mines & Mineral Resources, Memoir
27: 173 pp.
Sutherland, P. K., and Manger, W. L., 1984, Morrowan brachiopods in thetype "Derryan" Series (Pennsylvanian), New Mexico (abs.): GeologicalSociety of America, Abstracts with Programs, 16: 115.
Thompson, M. L., 1942, Pennsylvanian System in New Mexico: New
Mexico Bureau of Mines & Mineral Resources, Bulletin 17: 92 pp.

Morrowan—Atokan (Pennsylvanian) conodont biofacies, south-central New Mexico and westernmost Texas
David A. Kaiser and Walter L. Manger
Introduction
Environmental and microfacies control on theoccurrence, abundance, and distribution of conodontelements has been recognized for some time (Merrill andvon Bitter, 1976; Clark, 1984), and that collective influence
has become encompassed by the term biofacies. Biofacies
are recognized broadly in a variety of depositional settings,
but analysis of the lithologic successions preserving theconodonts and presumably influencing their occurrences
has not been rigorous (Driese et al., 1984, being a notable
exception). Three localities in southern New Mexico andwesternmost Texas (Fig. 1) were sampled for conodonts
and associated carbonate microfacies to quantify theconodont biofacies present. Those localities represent atransition from the Robledo shelf to Orogrande basin within
a carbonate-dominated regime of Morrowan and Atokanage. Lithostratigraphic nomenclature for the interval of
study is unsettled, but it would include the La Tuna andBerino members of the Magdalena Formation (Nelson,
1940) and the Derry Series of Thompson (1942).
Carbonate microfacies
Carbonate lithologies from the three sections were described by Clopine (1990), who recognized seven microfacies tracts that include a textural spectrum from mudstones
to grainstones, but are dominated by mud-bearing litholo-gies (Fig. 2). Other previous investigations have included
the Vinton Canyon section and exposures near the Bishop
Cap section and have provided similar interpretations of the
microfacies (Mims, 1971; Osleger, 1981). The present study
concentrated on Clopine's skeletal wackestone/packstone(Fig. 2-C/D) and crinoidal grainstone/packstone (Fig. 2-E)
microfacies. Modal analyses of 300 counts on 85 thin-sections cut from those lithologies further recognized a mixed
skeletal wackestone/packstone and intraclast packstone microfacies (Kaiser, 1990) (Fig. 2).
The depositional history of shallow-marine carbonates
on the Robledo shelf and in the Orogrande basin can beinterpreted by the standard facies scheme of Wilson (1975).
Transgression began in the Morrowan from the Orograndebasin and covered the Robledo shelf by upper Atokan time.
Subsidence in the basin kept pace with deposition retaining
the low-energy, shallow-marine character, particularly highmud content, and blanket distribution of the Pennsylvaniancarbonates throughout the region.
Conodont biofacies
Eight platform elements assignable to 24 form species
are represented among the 3318 elements recovered from76 productive samples; six samplcs were barren. Each sample was assigned to a carbonate microfacies and its
conodont occurrence data were tabulated. Biofacies werebased on (1) element abundances, (2) platform to ramiform
ratios (P/ R ratios), and (3) form-generic abundances (Fig.
3). Three conodont biofacies were identified: Adetognathus,
Idiognathodus, and Idiognathoides.
The Adetognathus biofacies is recognized by the total
dominance of that form taxon, which may be the only element recovered in a particular sample. The only other associated platforms are Rhachistognathus and Idiognathoides,
recovered from a total of only three samples. The biofacies isdeveloped in the shallow, near-shore lithologies of the mixed
skeletal wackestone/packstone and intraclast pack-192stone microfacies. Environments associated with the Adetognathus biofacies had moderate energy levels and possibly
fluctuating salinities. Most samples from this biofacies exhibit low abundance and low P/R ratios. High abundancesand high P/R ratios in samples from this biofacies represent
lag concentrates, while low abundances and high P/R ratios
probably represent transportation into otherwise inhospitableenvironments. This biofacies is equivalent to that previously
recognized by Merrill and von Bitter (1976), Driese et al.
(1984), and Davis and Webster (1985). It was confined torocks of Morrowan age, and reflects the transgressive natureof those strata across the Robledo shelf (Fig. 4). Its
dominance of the Morrowan section at Vinton Canyon reflects persistent shallow environments as deposition kept
pace with subsidence in the Orogrande basin.
The Idiognathodus biofacies (Figs. 5, 6) is recognized insamples in which that taxon comprises 50% or more of theplatform elements recovered. Associated platform elements
include: Neognathodus, Streptognathodus, Idiognathoides,
Adetognathus, and Diplognathodus. Three samples at
Bishop Cap fell below the abundance level for
Idiognathodus, but were included in this biofacies becauseof the similar elemental association. The Idiognathodus
biofacies is developed in the crinoidal packstone/grainstone
microfacies that reflects more off-shore, deeper, marine
environments of normal salinity basinward of the
Adetognathus biofacies. Samples from this biofacies
typically have high abundances but low P/R ratios,
indicating optimal living conditions (natural assemblages
suggest a P/R ratio of 0.2). Lag concentrates of highabundance and high P/R ratios also occur in samples fromthis biofacies. The Idiognathodus biofacies was recognized
by Merrill and von Bitter (1976) and Driese et al. (1984). It
is represented in the highest Morrowan section at Bishop
Cap and Vinton Canyon, and comprises the entire Atokanrecord at all three sections studied (Fig. 4). The replacementof the Adetognathus biofacies by the Idiognathodus
biofacies is consistent with the continued transgression anddeepening of environments across the Robledo shelf and in
the Orogrande basin during Atokan time.
An Idiognathoides biofacies (Fig. 5) is recognized insamples in which that taxon comprises 50% or more of theplatform elements present. Associated platform elements
include: Neognathodus, Adetognathus, Streptognathodus,
Idiognathodus, and Rhachistognathus. This biofacies wasfound in two samples of the mixed skeletal wackestone/packstone microfacies at Bishop Cap, and reflects a transition from the Adetognathus to Idiognathodus biofacies.
Too large a sampling interval may account for the absenceof this biofacies at Vinton Canyon, while transitional intervals may not have been preserved at the type Derryan
section. This biofacies is equivalent to that recognized by
Davis and Webster (1985), but may represent the Neognathodus biofacies of Merrill and von Bitter (1976) andDavis and Webster (1985) as well. Abundant Neognathodus

may reflect stressed environments allowing success of onlythat platform taxon.
Biostratigraphic considerations
The biofacies domination of conodonts exhibited by theMorrowan—Atokan succession limits their biostratigraphic
utility in the Orogrande basin and on the Robledo shelf. The
Morrowan zonal name-bearers are present, but their occurrences are sporadic and all may appear simultaneously ina lag concentrate (e.g. Bishop Cap). Most samples contain
Adetognathus to the exclusion of other platform taxa. Neither
the Morrowan—Atokan nor Atokan—Desmoinesian boundaries are accompanied by a pronounced change in conodonts,
although the abrupt appearance of Diplognathodus
coloradoensis suggests an unconformity at the former
boundary in the type Derryan and Bishop Cap sections.
Atokan Neognathodus occurrences offer general- discrimination of lower and upper portions of the series that arecompatible with associated foraminifers. However, no zonation is recognizable through that interval on the basis ofconodonts.
References
Clark, D. L. (ed.), 1984, Conodont biofacies and provincialism:
Geological Society of America, Special Paper 196, 340 pp.
Clopine, W. W., 1990, Middle Pennsylvanian fusulinid biostratigraphy of
the type Derryan area in southern New Mexico and westernmost
Texas: Unpublished PhD dissertation, University of Oklahoma,
Norman, 271 pp.
Davis, L. E., and Webster, G. D., 1985, Late Mississippian to early
Pennsylvanian conodont biofacies in central Montana: Lethaia, v. 18,
pp. 67-72.
Driese, S. G., Carr, T. R., and Clark, D. L., 1984, Quantitative analysis ofPennsylvanian shallow-water conodont biofacies, Utah and Colorado; in
Clark, D. L. (ed.), Conodont biofacies and provincialism: GeologicalSociety of America, Special Paper 196, pp. 233-250.
Kaiser, D. A., 1990, Conodont biofacies analysis of the Lower to Middle
Pennsylvanian (Morrowan and Atokan), south-central New Mexico andwesternmost Texas: Unpublished MS thesis, University of Oklahoma,
Norman, 115 pp.
Kottlowski, F. E., 1960, Summary of Pennsylvanian sections in southwestern New Mexico and southeastern Arizona: New Mexico Bureau of
Mines & Mineral Resources, Bulletin 66, 187 pp.
Merrill, G. K., and von Bitter, P. H., 1976, Revisions of conodont carbonate biofacies nomenclature and interpretations of environmental controls in Pennsylvanian rocks of eastern and central North America:
Royal Ontario Museum, Life Sciences Contributions, no. 108, 46 pp.
Mims, R. L., 1971, Microfacies analysis of the La Tuna Formation (Morrowan), Vinton Canyon, El Paso County, Texas: Permian Basin Section,
Society of Economic Paleontologists and Mineralogists, Guidebook to
the Robledo and Franklin Mountains, pp. 87-106.
Nelson, L. A., 1940, Paleozoic stratigraphy of Franklin Mountains, West
Texas: American Association of Petroleum Geologists, Bulletin, v. 24,
pp. 157-172.
Osleger, D. A., 1981, Stratigraphy and microfacies analysis of the La Tuna
Formation (Morrowan—Atokan), Franklin Mountains, Texas and New
Mexico, and Bishop Cap, New Mexico: Unpublished MS thesis, University of Texas at El Paso, 138 pp.
Thompson, M. L., 1942, Pennsylvanian System in New Mexico: New
Mexico Bureau of Mines & Mineral Resources, Bulletin 17, 92 pp.
Wilson, J. L. W., 1975, Carbonate facies in geologic history: SpringerVerlag, Berlin, 471 pp.

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