The Geology of Big Bend National Park

The Geology of Big Bend National Park:
What have we learned since Maxwell
and others (1967)?
Field Trip Guide
John C. White, editor
Field Trip Contributors:
Dan S. Barker, University of Texas at Austin
Tim W. Duex, University of Louisiana Lafayette
Richard J. Erdlac, Jr., University of Texas Permian Basin
Gary L. Kinsland, University of Louisiana Lafayette
Thomas M. Lehman, Texas Tech University
Rolfe D. Mandel, University of Kansas
Don F. Parker, Baylor University
Carol Purchase, National Park Service
Kevin M. Urbanczyk, Sul Ross State University
John C. White, Baylor University and Sul Ross State University
John Zak, Texas Tech University
THE GEOLOGY OF BIG BEND NATIONAL PARK: WHAT HAVE WE
LEARNED SINCE M AXWELL AND OTHERS (1967)?
Geological Society of America South-Central Meeting, April 2002,
Sul Ross State University, Alpine, Texas.
Table of Contents
Page
Introduction............................................................................................................1
Day 1.....................................................................................................................3
The Geology of Elephant Mountain (D.S. Barker).................................................5
Stop 1. Elephant Mountain WMA..........................................................................6
Late Quaternary alluvial stratigraphy in Calamity Creek valley, Elephant
Mountain Wildlife Management Area (R.D. Mandel)
Stop 2. West Entrance to Big Bend National Park..............................................12
The Terlingua Monocline (R.J. Erdlac, Jr.)
Old Maverick Road (T.M. Lehman)......................................................................15
Stop 3. Peña Mountain.......................................................................................15
Pen and Aguja Formations (T.M. Lehman)
Stop 4. Santa Elena Canyon...............................................................................21
Terlingua Creek Delta (K.M. Urbanczyk)
Stop 5. Horseshoe Canyon.................................................................................24
Horseshoe Canyon Volcanic Dome (D.F. Parker)
Stop 6. Tuff Canyon............................................................................................33
Tuff Canyon (D.S. Barker)
Stop 7. Sotol Vista..............................................................................................36
The Sierra Quemada Caldera (T.W. Duex and G.L. Kinsland)
Terlingua Monocline Overview (R.J. Erdlac, Jr.)
Day 2
Stop 8. The Basin...............................................................................................43
Extracaldera vents of the Pine Canyon Caldera (J.C. White)
Stop 9. Pine Canyon Vista..................................................................................47
Pine Canyon watershed research program (K.M. Urbanczyk and J. Zak)
The Pine Canyon Caldera (K.M. Urbanczyk and J.C. White)
Stop 10. Boquillas Overlook................................................................................50
Geomorphologic changes in the Rio Grande / Rio Bravo channel and
floodplain (C. Purchase)
Stop 11. Boquillas Canyon..................................................................................51
Rio Grande Holocene Valley Fill (R.D. Mandel)
Stop 12. Fossil Bone Exhibit...............................................................................54
A view of the Tornillo Group (T.M. Lehman)
Tornillo Flats Grasslands (C. Purchase)
References Cited.................................................................................................61
Cover Photo: Landsat image of Big Bend National Park.
INTRODUCTION
Big Bend National Park is often referred to as the only place in North
America where the Appalachia Mountains meet the Rocky Mountains, and this
statement does not contain as much hyperbole as one might think. Within the
northern part of the park, at Persimmon Gap, Ordovician to Pennsylvanian
pelagic and flysch strata deformed during the late Pennsylvanian MarathonOuachita orogeny are overlain by middle Cretaceous limestone, all of which was
deformed during the late Cretaceous-early Tertiary Laramide orogeny; in Big
Bend National Park, West Coast geology does indeed meet East Coast geology!
Even with this remarkable diversity concentrated in such a small part of the park,
the geology at Persimmon Gap does not completely represent what one can
expect to find in Big Bend National Park. In addition to Paleozoic deep marine
strata deformed during the Ouachita-Marathon orogeny and middle Cretaceous
shallow marine strata deformed during the Laramide orogeny, within the park one
can also find:
• A complete section of middle to late Cretaceous marine sediments
correlative with central Texas strata.
• A late Cretaceous to early Tertiary regressive sequence (shallow marine
to terrigenous) in which are found numerous important vertebrate fossils. Upper
Cretaceous strata have yielded numerous Mosasaurs, Quetzalcoatlus northropi
(the largest animal that ever flew), and many Sauropod bonebeds.
• Middle Tertiary (Eocene to Oligocene) mafic to felsic volcanism, erupted
from two calderas, numerous volcanic domes, and several unidentified vents.
• Many igneous intrusions, including laccoliths (Government Spring,
Grapevine Hills, Rosillos Mountains, McKinney Hills), and several dikes and
plutons related to identified volcanic centers.
• Miocene to Recent Basin and Range normal faults (a third orogeny)
responsible for many of the high, spectacular scarps found in the park (e.g.,
Mesa de Anguila – Sierra Ponce at the mouth of Santa Elena Canyon, the Sierra
del Carmen – Sierra del Caballo Muerto).
• Thick sequences of bolson-filling alluvium associated with the two major
grabens (the Tornillo graben on the east and the Castolon graben on the west)
and overlying Quaternary soils and pediments.
As one of the most geologically diverse National Parks in the United
States, Big Bend National Park has been the subject of much inquiry, beginning
with the early exploration by Robert T. Hill (1901, 1902) and Johan A. Udden
(1907). However, the first comprehensive geologic study and map of the park
came with the mapping by Ross A. Maxwell, John T. Lonsdale, Roy T. Hazzard,
and John A. Wilson intermittently between the years 1936 and 1963. The results
of these investigations were published by the University of Texas at Austin
Bureau of Economic Geology in 1967 (Maxwell and others, 1967), which remains
to this day the only complete geologic map of Big Bend National Park ever
published. Since 1967, many studies—too numerous to list here—have been
conducted in Big Bend National Park that have built on the pioneering work of
Maxwell, Lonsdale, Hazzard, and Wilson, that have greatly improved our
1
understanding of the geologic history of the park, especially in the fields of
stratigraphy, paleontology, volcanology, structural geology, geoarchaeology, and
geomorphology. In the past decade, much of the focus in the park has been on
Environmental Geology, as draught and dams endanger the Rio Grande / Rio
Bravo, and as the air quality deteriorates. On this field trip and in this guidebook,
we will attempt to present a summary of how our understanding of the geology of
Big Bend National Park has improved over the past three decades. This is
certainly not a topic that can be addressed adequately on a two-day field trip or in
a fifty-something page guidebook, but it will hopefully help inspire future studies
in Big Bend National Park—including the production of a revised map of the
park—that will continue to contribute to our growing knowledge of this varied and
unique area.
ACKNOWLEDGEMENTS
The editor would like to thank several people who have contributed to or assisted
in the production of this field trip and guidebook. This field trip (and
corresponding symposia) was conceived by Kevin Urbanczyk, who has a longterm commitment to furthering our knowledge of Big Bend geology. Kevin also
contributed stops for this trip, as have Dan Barker, Tim Duex, Richard Erdlac, Jr.,
Gary Kinsland, Tom Lehman, Rolfe Mandel, Don Parker, Carol Purchase, and
John Zak. These contributions are greatly valued, and without them this field trip
would not have been possible. Mary Kay Gordon, department secretary,
provided great assistance with working out field trip logistics. Michael O’Ferrall
assisted me in the field during the production of the road log. Ross Maxwell’s
successor, Frank Deckert, the current park superintendent, is also owed a debt
of gratitude for providing fee waivers for our group. Additionally, I extend thanks
to several Sul Ross students who have agreed to help during the field trip—I’d list
them by name, but as of this writing, I’m not sure who they are! Last, I would like
to thank Jennifer White, the person who edits the editor.
2
FIELD TRIP ROADLOG AND STOP ARTICLES
DAY ONE
This field trip begins at the Sul Ross State University Center parking lot and
departs at 7:00 am. Day one begins in Alpine and ends at the Longhorn Ranch
Motel after a day in the west side of Big Bend National Park (Figure 1).
Exit the campus through Entrance 4, turning left (south) on North Harrison St.
(Sate Highway (SH) 223, the shortest numbered highway in the state) and reset
your odometer to zero.
0.0
SRSU Entrance 4. Turn left on Harrison St.
Intersection of Harrison Road (SH 223) and East Avenue E (Hwy 90/67).
0.1
Turn right (west) and get into the left lane as quickly as possible.
0.3
Intersection of Avenue E (Hwy 90/67) and SH 118. Turn left (south).
11.5
Mount Ord (el. 6700 ft) is at 9:00.
14.2 Border Patrol checkpoint. Cathedral Mountain (el. 6860 ft) is at 3:00. The
base of Cathedral Mountain is mapped as Duff Formation tuffaceous
sedimentary rock and is overlain by the Mitchell Mesa Tuff, which forms the lower
bench. The Mitchell Mesa Tuff is the most widespread ignimbrite in the TransPecos and represents the caldera-forming eruption of the Chinati Mountains
caldera (~32 Ma), which lies ~50 miles southeast of Cathedral Mountain (Henry
and Price, 1984; Henry and McDowell, 1986). Overlaying the Mitchell Mesa Tuff
is the Tascotal Formation, another sequence of tuffaceous sedimentary rocks
that represent the volcaniclastic alluvial apron that ringed the Chinati Mountains
caldera (Walton, 1979). The Tascotal Formation is capped by basalt mapped by
McAnulty (1955) as Rawls Formation, but is almost certainly not. These units
represent the upper part of Goldich and Elms’ (1949) Buck Hill series (Table 1).
From this point, the highway descends through the Buck Hill series to the
underlying Gulfian (Cretaceous) rocks that crop out in the O2 flats.
15.9 Entrance to Woodward Ranch on the right side of the road. For a small
fee, visitors are allowed to camp at Woodward Ranch and collect agate and opal
that weathers out vesicles in the Sheep Canyon basalt.
20.2
The road passes through an outcrop of Cottonwood Springs Basalt.
21.3
Roadcuts through lava flows of the Potato Hill Andesite.
21.6
A flow of Sheep Canyon Basalt with basal autobreccia.
3
Figure 1. Big Bend National Park and surrounding area. Numbers refer to field trip stops; stop 1 at Elephant Mountain is not shown.
“Rawls” Basalt
Tascotal Formation
Mitchell Mesa Tuff
Duff Formation
Pruett
Formation
Cottonwood Springs Basalt
Sheep Canyon
Basalt
Potato Hill Andesite
Crossen Trachyte
Table 1. Buck Hill Series, adapted from Goldich and Elms (1949). South of
Kokernot Mesa, where the Cottonwood Springs Basalt does not crop out, the Pruett and
Duff Formations cannot be distinguished and are often mapped as either “Pruett/Duff
undifferentiated” or as the Devils Graveyard Formation (Stevens and others, 1984). This
predominantly tuffaceous sedimentary unit is stratigraphically equivalent to the Chisos
Formation in Big Bend National Park (Maxwell and Dietrich, 1970).
24.8 Eocene lacustrine limestone (Pruett Formation) overlain by a flow of
Sheep Canyon basalt. (Collinsworth and Rohr, 1986)
26.4 Stop 1. Turn left into the entrance for Elephant Mountain Wildlife
Management Area (WMA).
THE GEOLOGY OF ELEPHANT MOUNTAIN
Dan S. Barker, Department of Geological Sciences, University of Texas at
Austin.
Elephant Mountain (el. 6230 ft.) is capped by a discordant sheet (“trapdoor
intrusion”) of resistant phonolite and nepheline trachyte. This is the largest of at
least thirty such intrusive bodies defining a NNW belt at least 350 km long and
only 20 km wide (Potter, 1996).
Elephant Mountain lies within a few kilometers of the Ouachita Front, the
buried northwestern limit of Paleozoic thrusted metasedimentary rocks. The belt
of alkalic intrusions crosses this and the Grenville Front 160 km farther north,
without showing clear chemical and isotopic effects of basement variation
(James and Henry, 1993; Potter, 1996). The Elephant Mountain igneous rock
varies in texture and in amount of nepheline. Anorthoclase is the dominant
mineral; in many samples analcime replaces nepheline. The mafic minerals
include fayalite, sodic clinopyroxene and sodic amphibole. Chemical analyses of
Elephant Mountain samples and more than 1500 other Trans-Pecos igneous
rocks are available as PDF files at http://www.geo.utexas.edu/faculty/barker.
5
Stop 1. LATE QUATERNARY ALLUVIAL STRATIGRAPHY IN CALAMITY
CREEK VALLEY, ELEPHANT MOUNTAIN WILDLIFE MANAGEMENT AREA
Rolfe D. Mandel, Department of Geography, University of Kansas, Lawrence.
Stop 1 provides an opportunity to examine late Quaternary alluvium in
Calamity Creek valley in the area of Elephant Mountain. This locality was the
setting for Claude C. Albritton, Jr. and Kirk Bryan’s benchmark studies that linked
Quaternary landscape evolution to climatic change (Albritton and Bryan, 1939;
Bryan and Albritton, 1943). It also is the area where J. Charles Kelley, T. N.
Campbell, and Donald J. Lehmer conducted one of the earliest and most
significant geoarchaeological investigations in North America (Kelley and others,
1940).
Albritton and Bryan (1939) recognized three alluvial stratigraphic units in
Calamity Creek valley: the Neville, Calamity, and Kokernot formations, in order
from the oldest to youngest (Figure 2). They suggested that the Neville formation
aggraded during the late Pleistocene. This interpretation was based on the
presence of elephant and horse bones in the Neville alluvium, though the
integrity of the faunal remains was not considered. The Calamity formation is
inset into the Neville formation (Figure 2) and was considered post-Pleistocene in
age. According to Bryan and Albritton (1943), artifacts recovered from the upper
part of the Calamity formation indicated that most of this stratigraphic unit
aggraded before ca. A.D. 700 (1,300 yr B.P.). The Kokernot formation fills
channels cut into the Neville and Calamity formations and occurs as a thin
veneer (overbank deposit) across the valley floor (Figure 2). Bryan and Albritton
(1943) suggested that aggradation of the Kokernot formation was underway prior
to A.D. 1200 (800 yr B.P.) and continued until A.D. 1880-1890 (ca. 100 yr B.P.).
As with the Calamity formation, they used archaeological evidence to infer the
age of the Kokernot formation.
Based on their work in Calamity Creek valley, Bryan and Albritton (1943)
concluded that alternating stages of erosion and deposition in “valley flats” of the
Davis Mountains reflect fluctuations in climate, with alluviation and dissection
occurring during moist and dry periods, respectively. They also suggested that
the accumulation of caliche (calcium carbonate) in buried soils in Calamity Creek
valley is a product of relatively arid conditions, whereas the “humic” (organic-rich)
soils are products of former humid conditions. We now know that the driving
forces and temporal and spatial patterns of erosion and deposition in drainage
networks are much more complex than Bryan and Albritton’s simple model (see
Schumm, 1977; Mandel, 1994). However, their use of soil properties to infer past
climatic conditions remains valid.
The alluvial “formations” defined by Claude Albritton and Kirk Bryan are
exposed in a cutbank near the headquarters for the Elephant Mountain Wildlife
Management Area (Figure 3). Calamity Creek has migrated laterally into a low
terrace, exposing a 5.6-m-high section of valley fill. The lower half of the section
was described and sampled in 1996 (Mandel, 1996; Table 1). A unit of silty
alluvium (Kokernot formation) composes the upper 60 cm of the section. This unit
has been slightly modified by soil development and probably is less than 1,000
6
Figure 2.
7
years old.
The Kokernot formation overlies a moderately developed soil developed at
the top of a thick package of silty and loamy alluvium (Calamity formation). The
most striking feature of the Calamity formation is a thick, dark buried soil at a
depth of 267-467 cm below the surface of the terrace. The buried soil has a
strongly expressed Bk-BCk profile and is developed in fine-grained alluvium in
the lower half of the Calamity formation (Table 2 and Figure 3); the A horizon
was stripped off by erosion before burial. Humates from the upper 10 cm of
Bk1b2 horizon yielded a radiocarbon age of 6,360+70 yr B.P. (Tx-9018). Given
that it is a single radiocarbon date determined on soil, it must be viewed with
caution. Nevertheless, this date suggests that aggradation of the Calamity
formation was interrupted by landscape stability and soil development during the
early part of the middle Holocene.
An abrupt, wavy boundary separates the Calamity formation from a
remnant of a buried soil developed in fine-grained alluvium that quickly grades
downward to sand and gravel. This lowermost buried soil is represented by a
truncated Bk horizon with stage II carbonate morphology (common films, threads,
and nodules); hence, it fits the description of Albritton and Bryan’s (1939) Neville
formation.
In sum, Albritton and Bryan’s alluvial-stratigraphic framework for Calamity
Creek was remarkably accurate for its time. Although their alluvial “formations”
would be designated as allostratigraphic or lithostratigraphic units according to
modern stratigraphic nomenclature, Albritton and Bryan had a keen sense of the
complex alluvial stratigraphy in Calamity Creek valley. In addition, their age
estimates for the alluvial deposits appear to be fairly accurate even though they
are not based on radiocarbon ages.
Additional geomorphological research, including intensive radiocarbon
dating, will be conducted in Calamity Creek valley over the next several years.
Much of this research will be coordinated with ongoing archaeological
investigations in the area of Elephant Mountain. Robert Mallouf, Director of the
Center for Big Bend Studies at Sul Ross University, will discuss some of the
results of their archaeological surveys during the presentation at Stop 1.
Exit Elephant Mountain WMA and turn left, continuing south on SH 118. As you
exit the WMA, reset your odometer to zero.
Elephant Mountain WMA Entrance. Kokernot Mesa lies at 12:00; the
0.0
mesa is capped by Crossen Trachyte, which overlies undifferentiated Pruett
Formation tuffaceous sediments. These units represent the lowest within the
Buck Hill series, and within the Tertiary system in the southern Davis Mountains.
7.6
Road cuts through Boquillas Formation (Upper Cretaceous) flaggy
limestone.
11.9 Santiago Mountain at 9:00. Like Elephant Mountain, Santiago Mountain is
one of the several nepheline trachyte – phonolite intrusions described earlier.
8
____________________________________________________________________________
Table 2. Description of the lower half of section along Calamity Creek, Elephant Mountain Wildlife
Management Area, Stop 1.
______________________________________________________________________________
Landform: T-1 terrace
Slope: 1 percent
Drainage class: Well drained
Vegetation: Creosote bush, mesquite, prickly pear, bristlegrass, and fluff grass
Described by: Rolfe Mandel
Date described: November 16, 1996
Remarks: Humates from the upper 10 cm of Bk1b2 horizon yielded a radiocarbon age of 6,360+70 yr B.P.
(Tx-9018).
______________________________________________________________________________
Depth
Soil
(cm)
Horizon
Description
“CALAMITY FORMATION”
267-292
Bk1b2
60% brown (7.5YR 5/4) silty clay loam, brown (7.5YR 4/4) moist, 40%
brown (7.5YR 5/3) silty clay loam, brown (7.5YR 4/3) moist; weak medium
subangular blocky structure parting to weak fine subangular blocky; hard;
common fine films and threads of calcium carbonate; few fine flecks of
charcoal; strong effervescence; gradual smooth boundary.
292-332
Bk2b2
50% brown (7.5YR 5/4) silty clay loam, brown (7.5YR 4/4) moist, 25%
brown (7.5YR 5/3) silty clay loam, brown (7.5YR 4/3) moist, 25% brown
(7.5YR 4/2) silty clay loam, dark brown (7.5YR 3/2) moist; moderate
medium and coarse prismatic structure parting to moderate medium
subangular blocky; very hard; common pressure faces on peds; common fine
films and threads of calcium carbonate; strong effervescence; gradual
smooth boundary.
332-396
Bk3b2
Brown (7.5YR 5/3) silty clay loam, brown (7.5YR 4/3) to dark brown
(7.5YR 4/2) moist; few fine distinct yellowish red (5YR 4/6) mottles;
common cracks 3-5 mm wide filled with brown (7.5YR 5/4) silty clay loam,
brown (7.5YR 4/4) moist; moderate medium and coarse prismatic structure
parting to moderate medium subangular blocky; very hard; common
pressure faces on peds; common films and threads of calcium carbonate,
especially along root paths; strong effervescence; gradual smooth boundary.
396-457
BCkb2
Brown (7.5YR 5/3) silt loam, brown (7.5YR 4/3) to dark brown (7.5YR 4/2)
moist; common fine faint brown (7.5YR 5/4) mottles; weak fine and
medium angular blocky structure; very hard; common films and threads of
calcium carbonate, especially along root paths; few round pebbles; strong
effervescence; abrupt wavy boundary.
457-495
Bkb3
495-560+
C’
“NEVILLE FORMATION”
Brown (7.5YR 4/4) silty clay loam, dark brown (7.5YR 3/4) moist; common
fine faint brown (7.5YR 5/4) mottles; weak fine and medium prismatic
structure parting to weak fine subangular blocky; hard; many films and
threads and common fine hard nodules of calcium carbonate (stage II); few
round pebbles; strong effervescence; clear smooth boundary.
Stratified sand and gravel; single grain; loose; gravels are imbricated.
9
Figure 3.
10
17.6
Buck Hill, a quartz syenite sill, at 2:00.
18.6 First view of the Chisos Mountains, rising high in the background, with the
Christmas Mountains in the foreground.
25.3 One of several local bentonite mining companies in the area. The
bentonite is mined out of the Pruett Formation.
29.5
Nine-point Mesa, a 300 m thick quartz syenite sill, at 9:00.
31.8
Agua Fria Mountain, a peralkalic rhyolite laccolith, at 2:00.
36.7 Packsaddle Mountain at 3:00. The Christmas Mountains area is
characterized by numerous metaluminous to peralkalic, quartz trachyte to rhyolite
laccoliths that form domed mountains like Packsaddle, Agua Fria, and Hen Egg
Mountains (Henry and others, 1989). Here, the Packsaddle Mountain laccolith
has domed back Cretaceous-age Santa Elena Limestone.
The Terlingua – Christmas Mountains area was initially studied and
mapped my John T. Lonsdale (1940).
38.2 Longhorn Ranch Motel. We will return here at after we complete the stops
for Day 1.
41.9 Luna Vista Sill, a peralkalic quartz trachyte with high Na2O (~7.50 wt%)
and abundant arfvedsonite (~15 vol%) (Parker and others, 2000). The sill is
intruded into shale and siltstone of the Aguja Formation (Upper Cretaceous),
which we will discuss at Stop 3.
45.5 Wildhorse Mountain quartz syenite crops out along the east side of the
highway.
47.8 Willow Mountain, a quartz syenite intrusion with remarkably welldeveloped columnar jointing.
51.5 Study Butte (“STEW-dee B’yout”), Texas. Junction of FM 170 and SH
118. FM 170 is better known as the “Camino del Rio” and is consistently ranked
as one of the most scenic drives in Texas. Traveling west on FM 170, one
passes through Terlingua and Lajitas before paralleling the Rio Grande through
Big Bend Ranch State Natural Area and on to Presidio. Study Butte, like
neighboring Terlingua, was a major center of Cinnabar mining in the early part of
the 20th century (Ragsdale, 1976). Continue on SH 118 to Big Bend National
Park.
11
53.5 Stop 2. West entrance to Big Bend National Park. Park cars at the
entrance sign and prepare for a short hike. Remember that collecting rocks,
plants, artifacts, or any other samples without a permit in the park is illegal
– Leave your rock hammer in the vehicle!
Stop 2. THE TERLINGUA MONOCLINE
Richard J. Erdlac, Jr., Department of Geosciences, University of Texas Permian
Basin, Odessa.
The Terlingua monocline is at least 18 miles in length, with the western 11
miles having an overall strike of N70oW and 5 miles in the middle trending due
east. An additional 2 miles trends N65oE through this stop (Figure 4) along
Dawson Creek both west and east of U.S. 118. The monocline can be traced
eastward to the head of Dawson Creek within the Aguja and underlying Pen
Formation. It is highly eroded within the Pen clays and Aguja sandstones. It
may extend farther east but is obscured by erosion, intrusives, volcanic rocks,
and younger sediments.
Figure 4. This image is looking west along the dip slope of Terlingua Monocline in
Dawson Creek. This area is just within Big Bend Park boundary. Hogbacks composed
of Aguja Sandstone define this eastern part of the monocline. The overall strike of the
monocline at this location is N65oE. In the distance the flat horizon is composed of
Quaternary gravels that cover the dipping strata that define the monocline. The most
distant peaks on the right side of the image represent a small portion of the Terlingua
uplift that is named Reed Plateau.
The Terlingua monocline as a whole forms the northern boundary of Udden’s
‘Sunken Block’. At this stop (Figure 5) the ridge of the monocline displays a
sinusoidal shape along strike (Erdlac, 1988). Beds of Aguja define the monocline
12
with southerly dips between 13o to 57o. Dips within the Aguja Formation are
steeper along this ridge than in isolated outcrops of Aguja to the north, or in the
Javelina Formation to the south. Cross-bedding is readily observed within the
Aguja and paleosols, coincident with the Cretaceous/Tertiary boundary, have
been extensively investigated within the Javelina and overlying Black Peaks
Formation (Lehman, 1990). Small faults locally offset the strike of the monocline
(Figure 5). Although structural relief along the monocline has been estimated at
over 610 meters (Yates and Thompson, 1959), more recent field mapping
(Erdlac, 1988, 1990) suggests structural relief of 460 meters or less. The
monocline is cored by a high-angle reverse and strike-slip fault (north-side-up)
(Erdlac, 1988, 1990). Small out-of-the-syncline reverse or thrust faults are also
present in this segment and along other parts of the Terlingua monocline.
Figure 5. The Chisos Mountains form the backdrop for this image looking east along
the dip slope of the Terlingua Monocline in Dawson Creek. This area is just within Big
Bend Park boundary, with the location for this photo west of the actual Stop 2. Hogbacks
composed of Aguja Sandstone also define this part of the monocline. The monocline
turns from about N90oE to N65oE in trend. This image was taken looking east along the
N90oE segment of the monocline. A notch (right of center) along the monocline offsets
the Aguja ridge by 30 to 37 meters of left separation. No slickensides were found; thus
this offset could have a vertical component as well.
Westward, the Terlingua monocline forms the southern boundary of the
Terlingua uplift, a rhombic-shaped push-up feature that is both structurally and
topographically high (Erdlac, 1990). The monocline is here composed of steeply
dipping Santa Elena, Del Rio, Buda, and Boquillas Formations. Dip along this
part of the Terlingua monocline has increased, and ranges from 48o to 83o south.
This steep dip is especially evident in the area of Tres Cuevas Mountain (Figure
6). Croesus Canyon, which cuts through the monocline at this location from
13
north to south, provides the only access into the core of the monocline that well
displays the faults that control the location of the monocline. These faults display
slickenlines that demonstrate both the strike-slip and reverse-slip movement
involved with the formation of the monocline.
Figure 6. View east along steeply dipping Santa Elena Limestone from near the crest
of Tres Cuevas Mountain, the highest point along the Terlingua monocline. The Buda
and Del Rio Formations are low on the flank of the monocline. The Buda forms the
series of hogbacks along the right center part of the image. The Del Rio is located within
the topographic low between the Buda and the Santa Elena. The area within the
foreground and background of this image is cut by a deep canyon (Croesus Canyon) that
cuts through the monocline. The area in shadow on the left central side of the image
shows the eastern wall of this canyon.
Beyond its structural and tectonic importance to the region, the Terlingua
monocline and was the source of the second largest cinnabar deposits within the
continental United States (Yates and Thompson, 1959; Ragsdale, 1976).
Cinnabar was recovered from northeast-trending fractures and faults along the
crest of the monocline. The Del Rio appeared to form a prominent seal for
cinnabar emplacement. While local deposits of cinnabar are associated with
local intrusions, the greatest amount of cinnabar was recovered from mines
associated with the deep faulting controlling the placement of the monocline.
The deepest mine was the Chisos Mine at Terlingua, which reached depths of
840 feet. The owner of this mine, Howard Perry, would only allow Johan Udden
into the mine because Udden was able to understand the geological relations
involving cinnabar emplacement (Ragsdale, 1976). He was able to show the
miners where to dig in order to keep out of water wet fractures.
14
Upon returning to your vehicles, reset your odometer to 0.0.
0.0
Entrance to Big Bend National Park
1.0
Just past the Fee Station, turn right on to Old Maverick Road, an improved
dirt road.
OLD MAVERICK ROAD
Thomas M. Lehman, Department of Geosciences, Texas Tech University,
Lubbock.
The Old Maverick Road extends from the entrance station at the west end
of Big Bend Park, southwestward to the mouth of Santa Elena Canyon on the Rio
Grande. The road descends from the gravel-covered surface of a Quaternary
pediment onto exposures of the Upper Cretaceous Pen and Aguja Formations,
generally following the drainage of Alamo Creek, a tributary of the Rio Grande.
The higher peaks visible along the route of the Old Maverick Road (e.g.,
Rattlesnake Mountain, Pena Mountain) are capped by middle Tertiary sills that
have intruded the Upper Cretaceous strata. The best exposures of the Upper
Cretaceous strata are found around the margins of the intrusions.
2.2
Tule Mountain at 9:00 is capped by the Tule Mountain Trachyandesite of
the Chisos Group.
7.1
Chimneys Trailhead West. Peña Mountain at 12:00.
7.4
Luna’s Jacal
8.6
Stop 3. Pull completely off the road!
Stop 3. PEN AND AGUJA FORMATIONS
Thomas M. Lehman, Department of Geosciences, Texas Tech University,
Lubbock.
A Review of Late Cretaceous Sedimentation in the Big Bend region
Upper Cretaceous sedimentary rocks in the Big Bend region comprise a
northeastwardly thinning wedge of marine, paralic, and continental strata.
Exposures of these strata to the west and northwest in Chihuahua and in
Presidio County, are assigned to the Ojinaga, San Carlos, and El Picacho
Formations. Correlative strata in Big Bend, and in nearby Coahuila, are referred
to the Boquillas, Pen, Aguja, and Javelina Formations (Lehman, 1985).
Throughout Late Cretaceous time, a subduction zone existed along the west
coast of Mexico, with accretionary melange and flysch accumulating in an
offshore trench complex and forearc basin in what is now Baja California.
15
Figure 7. Correlation of Upper Cretaceous sedimentary rocks in Trans-Pecos
Texas and adjacent Mexico (from Lehman, 1986).
Magmatism in Sonora during Late Cretaceous time formed an extensive
subduction-related volcanic arc in the Sierra Madre Occidental. This volcanic
highland created the primary source terrain for much of the Late Cretaceous
sedimentation in the Big Bend region, as well as in the Parras and La Popa
basins along the Gulf Coast in Mexico. The continental basement upon which
the volcanic arc was built, was separated from the Big Bend region throughout
most of Cretaceous time by a subsiding deep-water marine basin - the
Chihuahua Trough. A shallow-water carbonate platform - the Coahuila Platform existed northeast of the Chihuahua Trough and formed the basement of the Big
Bend region. Subsidence of the Chihuahua Trough continued to exert a strong
influence on sedimentation well into Santonian time. Pelagic open marine
limestone of the Boquillas Formation accumulated on the platform while deep
marine black shale of the Ojinaga Formation accumulated in the adjacent trough.
As the Chihuahua Trough filled, and subsidence relative to the Coahuila Platform
slowed, clastic sediment gradually inundated the platform, and during
Campanian time the shoreline prograded northeastwardly across the Big Bend
region. Fluvial-dominated deltaic, sandy strandplain, coastal marsh, and swamp
deposits accumulated at that time to form the lower parts of the San Carlos and
Aguja Formations. The strandline reached a point near the present SantiagoSierra del Carmen range by Early Campanian time.
16
Figure 8. Late Cretaceous paleoenvironmental reconstructions of the Big Bend
region (modified from Lehman, 1985); A) Cenomanian through Santonian time, showing
deposition of Ojinaga Formation (Ko) in the Chihuahua Trough and Boquillas Formation
(Kb) on the Coahuila Platform, B) early Campanian time, showing coastal progradation
during deposition of the basal sandstone and lower shale of the Aguja Formation (Kag);
C) middle Campanian time, showing transgressive deposition of the Rattlesnake Mt and
McKinney Springs members; D) Maastrichtian time, showing deposition of the Javelina
Formation during onset of Laramide tectonism and formation of the Tornillo Basin.
17
Renewed marine transgression is recorded by an extensive tongue of
marine shale within the middle part of both the San Carlos and Aguja
Formations. Ammonites, oysters, and inoceramid bivalves from these deposits
indicate that this transgression occurred in Middle Campanian time.
Sedimentation occurred in sandy shoals, barrier islands, and coastal estuaries
that formed during this transgression. Southwestward thinning of the marine
shale tongue within the Aguja Formation indicates that the Middle Campanian
transgression did not extend inland much further than the present eastern margin
of the Chihuahua Tectonic Belt. During Late Campanian time, the strandline
once again prograded northeastwardly across the Big Bend region, depositing
the upper parts of the San Carlos and Aguja Formations.
With the onset of Laramide tectonism in Maastrichtian time, a dramatic
change in sedimentation occurred in the Big Bend region, a change recorded by
the presence of coarse extra-basinal detritus in the sediments, and a shift to
southeastward paleocurrent orientation in the El Picacho and Javelina
Formations. These strata of latest Cretaceous age are entirely fluvial in origin.
To the west of the Big Bend region, strata within the Chihuahua Trough were
folded and thrust eastward along the edge of the Coahuila Platform, as they
moved along a decollement surface initiated within evaporites deep in the basin.
This deformation culminated in formation of the Chihuahua Tectonic Belt - a
southern extension of the Rocky Mountain Cordilleran Fold and Thrust Belt. To
the east of the Big Bend region, a west-facing thrusted monoclinal uplift was
initiated along the present site of the Del Norte - Santiago - del Carmen range, a
feature similar to the classical Laramide basement-cored uplifts of the central
Rocky Mountains. During Maastrichtian and subsequent Paleogene time,
sedimentation was largely restricted to the area (the "Tornillo Basin") between
these two deformed regions (Lehman, 1991).
Local Upper Cretaceous Stratigraphy
The Pen and Aguja Formations weather recessively to form badlands
flanking the intermittent stream valleys and surrounding the eroded intrusive
rocks in the Alamo Creek and Terlingua Creek drainages. These strata are
exposed in a broad southeast-plunging anticline offset by a series of parallel
normal faults (generally down to the northeast) that comprise the Terlingua Abaja
Fault Group of Maxwell and others (1967). The Upper Cretaceous strata
comprise an intertonguing series of marine shales, deltaic sandstones, and
continental mudstones (Figure 8). Exposures along the flank of Rattlesnake
Mountain were designated the type area for what Udden (1907) called the
"Rattlesnake Beds". Adkins (1933) later renamed these beds the Aguja
Formation (Udden's term was already in use elsewhere) and established the type
area at Sierra Aguja ("needle peak") about 4 miles to the southwest. Lehman
(1985) later subdivided the Aguja Formation into several informal members that
reflect the intertonguing lithologies.
18
The broad flat area traversed by the Old Maverick Road is underlain
primarily by smectitic marine shales of the Pen Formation. These strata are
subject to dramatic shrink-and-swell action, resulting in "popcorn" weathering,
and a deep yellow limonitic weathering zone with surficial gypsum crystal crusts
due to decomposition of interstitial pyrite. The sediments contain biostromes of
the oyster Exogyra and rudist Durania, and scattered ammonites. The Pen
Figure 9. Stratigraphic relationships of members of the Aguja Formation and Pen
Formation in Big Bend National Park.
Formation is about 200 m thick in this area, and ranges from late Santonian to
early Campanian in age. Thin sandstones in the top of the Pen Formation
contain storm-generated shell beds with a diverse shallow marine molluscan
fauna dominated by bivalves and gastropods.
The Aguja Formation records two progradational cycles, the lower of
which is present only in the western part of the Big Bend region. The basal
sandstone member (10 m thick) is a progradational shoreface and deltaic
sandstone containing ammonites and inoceramid bivalves of late early
Campanian age. It is overlain by the lower shale member (70 m thick) which
consists of interbedded carbonaceous shale and lignite that accumulated in
coastal marshes and swamps landward of the shoreline (Record and Lehman,
1989). The lignitic shales are mined locally and sporadically as a source of soilconditioning "humates". In places, thin seams of bituminous-grade coal are
found that have also been mined, particularly in areas adjacent to the intrusive
rocks.
The lower shale member is overlain by a highly fossiliferous transgressive
marine sandstone - the Rattlesnake Mountain sandstone member of the Aguja
(Macon, 1994). This unit (10 m thick) contains abundant oysters (Flemingostrea
and Crassostrea) as well as inoceramids and ammonites indicative of middle
19
Campanian age. Shark teeth are also common in this deposit. Recently, the
remains of a gigantic sea turtle were also recovered from this unit.
A thin marine shale overlies the Rattlesnake Mountain sandstone. This marine
shale thickens eastward where it is referred to as the McKinney Springs tongue
of the Pen Formation (Mosley, 1992). Exposures along the Old Maverick Road
are near the landward pinchout of this shale, and it is here very thin (12 m) and
lignitic, suggesting accumulation in a lagoonal setting.
Overlying the McKinney Springs marine shale tongue is the Terlingua
Creek sandstone member, the second progradational deltaic/shoreface
sandstone of the Aguja. This unit is extensive over the entire Big Bend region,
and thickens dramatically (up to 30 m) in proximity to major deltaic distributary
channel complexes. It is relatively thin (8 m) in interdistributary areas.
Paleocurrent data indicate that the deltas which deposited the Terlingua Creek
sandstone prograded northeastwardly across the Big Bend region. The
paleoshoreline trended roughly northwest to southeast.
The Terlingua Creek sandstone is overlain by the upper shale member of
the Aguja, which was deposited primarily in fluvial coastal plain environments.
The upper shale member approaches 200 m in thickness in exposures along the
Old Maverick Road, but thins dramatically to the northeast, where it is typically
less than 100 m thick. Most of the terrestrial vertebrate fauna of the Aguja
Formation has been collected from sites in the upper shale member, and
indicates a Late Campanian to Early Maastrichtian age for this part of the
formation. The hadrosaurian ("duck-billed") dinosaur Kritosaurus and
ceratopsian ("horned") dinosaur Chasmosaurus dominate the fauna (Davies and
Lehman, 1989; Lehman, 1989; Wagner and Lehman, 2001). A variety of other
dinosaurs, including tyrannosaurids, ankylosaurs ("armored" dinosaurs),
pachycephalosaurs ("dome-headed" dinosaurs), ornithomimids ("ostrich"
dinosaurs) have also been found (Rowe and others, 1992; Lehman, 1997, 2001).
The giant crocodile Deinosuchus was also a prominent member of this fauna
(e.g., Anglen and Lehman, 2000), which inhabited open marsh habitats in what
was otherwise a closed-canopy tropical evergreen forest (Wheeler and Lehman,
2000; Lehman and Wheeler, 2001).
Note: A geologic map of the Old Maverick Road area, between Pena Mt and
Rattlesnake Mt, showing exposures of the Pen and Aguja Formations (from
Lehman, 1985) is included as an appendix.
13.5 The end of Old Maverick Road. Turn right (west) on the paved road (Ross
Maxwell Scenic Drive).
14.1 Stop 4. Santa Elena Canyon (Figure 9). Park your vehicle at the cul-desac parking lot and prepare for a very short walk to the mouth of Santa Elena
Canyon and the Terlingua Creek Delta.
From the Santa Elena Canyon parking lot, we return the way we came, traveling
east on Ross Maxwell Scenic Drive.
20
15.0
Santa Elena Canyon overlook.
16.2
Santa Elena Canyon take-out on right.
22.0
Cottonwood Campground on right.
Figure 10. Santa Elena Canyon. Strata from bottom to top are: Ktc, Telephone Canyon
Formation; Kdc, Del Carmen Limestone; Ksp, Sue Peaks Formation; Kse, Santa Elena
Limestone.
Stop 4. TERLINGUA CREEK DELTA
Kevin Urbanczyk, Department of Earth and Physical Sciences, Sul Ross State
University, Alpine, Texas.
We have conducted periodic surveys of the confluence of Terlingua creek
and the Rio Grande. The point of this work has been to get undergraduate
students involved in field research, and to document the characteristics of the
creek-river confluence. This is particularly interesting considering the chronic low
flow of the Rio Grande associated with drought in the 1990’s and early 2000’s
and increased upstream diversion.
The initial survey work was completed on 4-3-99 with the help of students
from the Introductory Geology (1401), Geology of West Texas (3301) and
Historical Geology (1402) classes from SRSU with supervision by Dr. Kevin
Urbanczyk and Dr. Jim Whitford-Stark.
21
The work consisted of setting up a TOPCON laser theodolite on a
limestone outcrop along the trail into Santa Elena canyon. The site was chosen
as to provide moderately easy accessibility and a clear view of most of the creek
channel (Figure 11). After the instrument was set up, two students with portable
radios and reflecting prisms were sent to numerous locations to survey the river
edge, the most recent creek channel locations and various vegetation sites. One
student was required to assist Dr. Urbanczyk at the instrument to take careful
field notes. Approximately 150 points were shot during a 3 hour period.
Figure 11. Terlingua Creek survey crew with the terrace and delta in front view.
Since 1990, the east margin of the Terlingua creek channel has moved
eastward about 40 meters at the expense of a terrace that had been covered by
vegetation. The west margin has also moved eastward about the same distance,
with the development of a new terrace covered with vegetation. Figure 11 shows
this terrace in front of the survey crew. Compare the vegetation seen below the
instrument in Figure 10 to the west side of the creek on DOQ shown in Figure 12.
Additionally, the creek flooding has pushed the margin of the Rio Grande
southward, and significantly narrowed its channel (see Figure 12). The amount
of Terlingua creek sediment along the northern banks of the Rio Grande indicate
that the most recent flooding occurred in creek, not in the Rio Grande.
22.6
Castolon village and store. LUNCH STOP.
23.7
Cerro Castellan (Figure 13) at 12:00. Cerro Castellan, like Horseshoe
Canyon (Stop 5) is a volcanic dome and one of several vents for the Wasp
Springs Formation and Burro Mesa Rhyolite.
22
25.2
Western end of the River Road.
25.4 Tuff Canyon. Vehicles will need to park here for the next stop after drivers
drop off passengers ~0.9 miles down the road.
Figure 12. 1990 DOQ of the Terlingua Creek delta with the results of the 1999 survey
superimposed (lower image). The green dot in the upper image shows the location of the
theodolite.
23
Figure 13. Cerro Castellan. Tcbm, Bee Mountain Basalt; Tcu, Chisos undifferentiated;
Tws, Wasp Springs Member; Tbmr, Burro Mesa Rhyolite.
26.3 Drop-off point for Stop 5. Drivers will need to drop off passengers here,
turn around, and park at Tuff Canyon (Mile 27.3). We will shuttle drivers from
Tuff Canyon back up to this point.
Stop 5. HORSESHOE CANYON VOLCANIC DOME
Don F. Parker, Department of Geology, Baylor University, Waco, Texas.
The surge deposits and nonwelded pyroclastic flows so excellently
displayed in the walls of Tuff Canyon suggest a nearby source (see previous
stop, this guidebook; Barker, 2000; Parker and others, 2000). Many of these
pyroclastic flows were erupted from the Horseshoe Canyon volcanic dome, the
focus of this stop (Figure 14). Horseshoe Canyon Dome is one of a half dozen or
so identified vent areas for Burro Mesa Rhyolite lava located along Ross Maxwell
Drive in a 16 km-long belt (Figure 15) (Henry and others, 1989; Holt, 1998).
Other eruptive vents are located along Burro Mesa (three or four vents), Kit
Mountain, Goat Mountain, and Cerro Castellan. Undoubtedly, others will be
discovered. The name Burro Mesa Rhyolite was first used by Maxwell and
others (1967), who correlated outcrops in the High Chisos Mountains with
exposures along Burro Mesa. I utilize the term “Burro Mesa Rhyolite” in a limited
24
sense in applying it to rocks of similar mineralogy, chemistry and age (~29 Ma;
Copeland and others, 1992) in the Cerro Castellan quadrangle and surrounding
areas.
Figure 14. View of Horseshoe Canyon Dome from Tuff Canyon parking lot. The
rhyolite dome is visible behind the eroded tuff ring.
Burro Mesa Rhyolite, in the strict sense, occurs in two principal types in
the Ross Maxwell Drive belt: sparsely-porphyritic rhyolite and abundantlyporphyritic rhyolite (Becker, 1976; Holt, 1998). The abundantly porphyritic
rhyolite appears to be limited to several vents located along Burro Mesa, where it
overlies sparsely-porphyritic rhyolite (Henry and others, 1989; Holt, 1998). The
upper member may have been derived from the lower member by about 15
weight percent fractionation of the observed phenocrysts, mostly anorthoclase
alkali feldspar. Rare mafic enclaves of trachyte found locally in the sparselyporphyritic rhyolite suggest that less evolved magma occurred deeper within the
magmatic system.
Each Burro Mesa vent was characterized by initial exposive eruptions that
formed a low-angle cone of non-welded surge deposits interbedded with nonwelded pyroclastic flow deposits. Lava, erupted from feeder dikes, then erupted
and may be observed at several centers where it cut upward through the
underlying pyroclastic cone (cf. Cerro Castellan; Burro Mesa Pouroff). Locally,
underlying pyroclastic rocks were welded from heat of the overlying lava.
25
Northwest Entrance, Horseshoe Canyon.
The Horseshoe Canyon area exposes an approximately 1.6 km diameter
volcanic dome and its associated, more extensive tuff cone (Figures 16 and 17).
Figure 15. Outcrop map of Burro Mesa Rhyolite along Ross Maxwell Drive.
After Holt (1998).
Our Stop begins along Maxwell Drive near milepost 19. We will hike southeast
about 2000 feet to the northwest entrance to Horseshoe Canyon while our
drivers return the vehicles to the Tuff Canyon parking area (there is no parking
available where we must leave the highway). At the entrance to the canyon,
non-welded surge and normal pyroclastic flows are overlain by a toe of Burro
Mesa lava (Figure 18). The immediate section of pyroclastic rocks beneath the
lava consists a basal pyroclastic flow, overlain by at least three sets of surge
deposits (Figures 19 and 20), which are in turn overlain by five non-welded
pumice-rich pyroclastic flows (each flow 1 meter or less in thickness), about 1
meter of pink surge deposits, and finally the brecciated base of the lava proper
(Figure 21). A steep exposure of cone deposits on the north side of the canyon
26
entrance appears to contain debris flow deposits with angular blocks of trachyte
up to 1 meter in diameter.
Figure 16. Geologic map of Horseshoe Canyon area. Stop begins as indicated
along Ross Maxwell Drive. Units keyed same as Figure 15, except Tbmdt = Burro Mesa
“Dam” Tuff and Qal = alluvium.
Figure 17. Cross section A-A’ from Figure 16. Tml = Mafic lava unit; Tt = Tuff
Cone; Tbml = Burro Mesa Rhyolite lava; Tg = older gravels.
The lava itself is nearly horizontal near the entrance to the canyon, but
steepens to the southeast. Flow banding is contorted although nearly parallel to
27
the base at the bottom of the flow, and steepens upwards into ramp structures
(Figure 22). Not visible at the stop, a brecciated zone occurs near the “U” of
Horseshoe Canyon, and appears to occupy a crevice-like zone extending
vertically through the unit. The north side of the canyon exhibits large, concentric
joints that parallel the front of the lava dome (Figure 23).
Figure 18. Base of Burro Mesa lava and underlying non-welded pyroclastic
deposits, west side of northwest entrance to Horseshoe Canyon.
Figure 19. Detail of surge sets from Figure 18.
28
Figure 20. Detail of pumice lapilli from non-welded pyroclastic flows. Knife is 8
cm long.
Figure 21. Brecciated base of Burro Mesa lava. Same locality as Figures 18, 19,
and 20. Cactus is 1 m high.
29
Figure 22. Large ramp structure in Burro Mesa lava at bend of Horseshoe
Canyon.
Figure 23. Concentric joints parallel to dome front, east side of northwest
entrance to Horseshoe Canyon. Joints cut flow banding. Note ramps in upper lava.
30
The vent area for the lava is not precisely located, but appears to have
been about 1000 feet (~300 m) south of the northwest entrance to Horseshoe
Canyon, judging from strike and dip of foliation in the lava (Figure 16). The
exposed feeder dike for the Cerro Castellan dome is only a few meters thick; a
similar presumed feeder dike for an abundantly-porphyritic dome and flow on the
northwest end of Burro Mesa is also very thin. These data suggest that a feeder
for Horseshoe Canyon might be very small in relation to the size of the dome.
The Southwestern Flank
Although we will not have time to visit it on this trip, the southwestern flank
of Horseshoe Dome exposes some important relationships (Figure 16). Normal
faulting formed a graben extending northwest through the map area, downdropping Tertiary gravels and isolating a segment of the dome. In this segment,
a thin edge of lava is overlain by densely-welded ignimbrite (Figures 24 and 25).
The ignimbrite is about 10 meters thick and appears to be of Burro Mesa
lithology, although detailed study has not been completed. Southwest of the map
area, the ignimbrite forms a small mesa, but, beyond that, its areal distribution
and source are unknown at present. I have informally named this unit the “dam
tuff,” from exposures along an eroded dam southwest of the mapped area.
Figure 24. Southwest flank of Horseshoe Dome. Densely-welded ignimbrite
overlies Burro Mesa lava with ramp structures
The western part of the mapped area is underlain by exposures of the
eroded tuff cone, which consists mostly of surge deposits and non-welded
pyroclastic flows containing lithic blocks up to 1 meter diameter (mostly of mafic
trachyte). The cone unit (Tt on Figure 16) overlies mafic trachyte lava (Tml) that
was mapped by Maxwell and others (1967) as Bee Mountain Basalt, and
appears to be similar to the mafic lava exposed in Tuff Canyon. The mafic
31
trachyte has been faulted against the Tt cone unit and the ignimbrite by a normal
fault with about 40 meters of vertical displacement (Figure 16).
Figure 25. Detail of densely-welded ignimbrite with lithic inclusions of
scoriaceous mafic trachyte
Suggested Work and Recommendations
The four field days spent mapping this small area reveal how much
remains to be discovered regarding the Cerro Castellan quadrangle and has
raised as many questions as it has answered. The eruptive center of Horseshoe
Dome appears documented, but where was the source of the Dam tuff? Also:
Maxwell’s mapping shows extensive Burro Mesa southeast of the map area. Is
this more lava from Horseshoe, another dome, or more ignimbrite? What are the
chemical and petrographic characteristics of the ignimbrite, and what is its
relation, if any, to the Burro Mesa lava? Obviously, more mapping and
petrologic study are warranted.
In regard to park development, a trail might be developed into Horseshoe
Canyon and related to the already-developed Tuff Canyon area. As such it
would complete the volcanological story already begun at Tuff Canyon, and
illustrated in the beautiful exposure of a similar vent at Cerro Castellan.
32
Stop 6. TUFF CANYON
Dan S. Barker, Department of Geological Sciences, the University of Texas at
Austin.
[A geologic map, sections, and photographs are in a guidebook for
nongeologists visiting Tuff Cayon (Barker, 2000)]. At this stop, we see distal
pyroclastic deposits from the Horseshoe Canyon vent area (the previous
stop) and a more complete section of the underlying mafic lava. The
intermittent stream of Blue Creek comes down from the western flank of the
Chisos Mountains and eventually enters the Rio Grande. Apparently
abandoning the segment of its old course through Horseshoe Canyon, Blue
Creek has vigorously incised Tuff Canyon, which is visible to the north from
the parking area.
The most obvious feature in this first view is a cross-canyon striking
NNW. Its western wall is a fault-line scarp parallel to a normal fault (down to
the east). The National Park Service has masterfully located three
observation platforms on the rim of Tuff Canyon. First we will walk to the
easternmost platform, across a gravel-armored surface.that impedes lateral
growth of the canyon.
Looking down from this observation platform, the farthest one
upstream, we can see a smoothly stripped dark surface on lava. This
surface begins to show more relief a little farther downstream. The lava is
probably part of the Bee Mountain “Basalt”, a 34.5 Ma package of flows with
varied compositions. Overlying the dark lava and forming the dry waterfall
visible upstream are paler layers of 29 Ma pyroclastic deposits that were
likely erupted from the vent(s?) at the previous stop. Note, in the top
surface of the dark lava, a network of light-colored cracks. Some of these,
at least, appear to be polygonal cooling joints that became filled with
pyroclastic debris that fell on the already weathered lava surface.
Now we retrace our steps, then past the parking area, and take the
short trail to the mouth of Tuff Canyon. Once inside the canyon, we will walk
upstream as far as the lava to examine its features, then turn around and
walk downstream but upsection, thanks to the westerly dip of the pyroclastic
units and their repetition by normal faults.
Below the observation platform we just visited, the dark lava forms a
riser and step in the canyon floor. A 10-meter section of lava has been
brought up on the eastern block of a normal fault. Among features visible
here are a large block with a pahoehoe surface and stretched vesicles, and
some not entirely convincing lava pillows. The pillows are just above the
contact between lava and the underlying hyaloclastite (“broken glass rock”).
The hyaloclastite is a matrix-supported breccia containing dark clasts of
chilled lava in a lighter colored matrix of clay, carbonate, and zeolites. It is
the product of thermal spalling and steam explosions generated when lava
entered standing water or water-saturated sediment. Perhaps the water had
been ponded by another lava flow. The pillows formed as lobes of lava
were quenched to thin glassy skins enclosing liquid lava. The upper,
33
massive, lava had less opportunity to interact with water because the
depositional surface had been built up by earlier lava. Here in one small
part of Big Bend we can see the two most common volcanic features on
Earth; hyaloclastites and lava flows cover most of the ocean floor.
Proceeding downstream we can get close views of the pyroclastic
units in the canyon walls. First, note that Tuff Canyon is misnamed for two
reasons. The deposits are not lithified, and are dominated by pumice lapilli
(pea-sized fragments of frothy rhyolite). “Tuff” refers to finer-grained and
consolidated deposits, but it is too late to change the Park signs. Also note
that the pumice fragments have a “woody” texture caused by elongated
vesicles. If the elongation had been caused by extrusion through a narrow
opening, each vesicle should have an oval cross section. These have
circular sections, giving rise to an interpretation (not believed by all
volcanologists) that such tubular pumice forms by enormous accelerations
during eruption, literally pulling the froth like taffy.
Some of the pyroclastic units show faint stratification (mostly by
particle-size variations), but also contain blocks of the dark lava. Certainly
the lava blocks and the pumice lapilli were far from being hydrologic
equivalents, and could not have been deposited from the same current in
running water. Instead, these weakly bedded units are interpreted as
deposits from pyroclastic surges, turbulent clouds of solid particles
suspended in gas. These surges moved along the ground surface at high
velocities. Locally, you can also see ballistic blocks of the mafic lava, with
impact sags where the falling blocks deformed the underlying layers. The
blocks of long-cold lava were probably thrown from a vent by phreatic
explosions when rhyolitic magma encountered groundwater. Although no
statistical study has been done, the size range of the blocks suggests that
the vent was less than 1 km away, and the asymmetry of the impact sags
indicates that the ballistic blocks came in from the east or northeast. Some
of the lava chunks are rounded (by weathering?) but are still called blocks,
not bombs. The term “bomb” implies that lava was still partially molten when
it was blown out.
Other pyroclastic units are more massive beds with tops defined by
coarser pumice. These are pyroclastic flow deposits, carried in by gas
clouds with higher concentrations of particles than surges, and probably
dominated by laminar flow. In contrast to the pyroclastic deposits around
the vent at the previous stop, these are not welded, apparently because
they had lost heat and were also thinner here. Some pyroclastic flow units
are cut by vertical gas-escape pipes. The ascending gas removed finer
particles from the pipes, leaving them loosely filled with a coarser lag
concentrate. In places, pipes cut upward through more than one flow unit,
demonstrating that the upper unit was deposited when the lower unit was
still losing its gases. Deposition of the pyroclastic units may have taken
only hours or days.
East-west channels filled by debris flow deposits locally interrupt the
pyroclastic section. As you walk back toward the canyon mouth, note the
34
“draperies” on the upper canyon walls. These are examples of case
hardening, where silica was dissolved from glass in the pumice and then
redeposit when water evaporated. Case hardening is probably a factor in
the preservation of potholes in the canyon walls.
27.3 Leave Tuff Canyon.
31.6
Mule Ears Overlook. The Mule Ears are intrusions of peralkalic rhyolite.
32.3 Goat Mountain (Figure 26). Like Horseshoe canyon and Cerro Castellan,
Goat Mountain is a vent area for the Burro Mesa Rhyolite. This interpretation of
Goat Mountain represents a very good example of “What we’ve learned since
Maxwell and others (1967)”: the exhibit for Goat Mountain gives Maxwell and
other’s interpretation—that this mountain represents a pale valley cut into older
Chisos units, which was then filled by an eruption of Wasp Spring “Flow-Breccia”
from the Pine Canyon volcanic center that was then incised into another valley,
which was filled by flows of Burro Mesa Rhyolite that also erupted from the Pine
Canyon center.
Figure 26. Goat Mountain volcanic dome.
34.0
Kit Mountain, another Burro Mesa Rhyolite vent area, at 9:00.
34.3
Eastern trailhead for the Chimneys Trail. The Chimneys are the eroded
remnants of Wasp Spring surge deposits erupted from the Kit Mountain
vent (Holt, 1998).
35.5
Burro Mesa Junction. Burro Mesa (Figure 27) is the type locality for Burro
Mesa Rhyolite. At least three vents areas for the Burro Mesa Rhyolite
have been identified in this area (Holt, 1998).
35
Figure 27. Burro Mesa volcanic dome.
38.7
Sotol Vista Overlook. Turn Right.
39.1 Sotol Vista parking lot.
Stop 7. SOTOL VISTA
Stop 7a. THE SIERRA QUEMADA CALDERA
Tim W. Duex and Gary L. Kinsland, Department of Geology, University of
Louisiana, Lafayette.
Looking southeast from Sotol Vista, many of the hilltops seen on the
skyline are related to the Sierra Quemada Caldera although it is difficult to
delineate a straightforward relationship because of the complex nature of the
caldera, the effect of Basin-and Range faulting, and differential erosion. Many
other topographic features in the area contain the eruptive products of the
caldera, most notably the Mule Ear Spring Tuff Member of the Chisos Formation.
The caldera is located about 4 mi (6 km) southeast of Sotol Vista and about 2 mi
(3 km) south of the South Rim of the High Chisos Mountains (Figures 28 and 29).
The Sierra Quemada is a significant event in the volcanic history of Big Bend
National Park because it produced a variety of igneous materials including the
oldest major felsic activity in the region. The materials will be discussed as those
that formed before, after, or during, collapse of the caldera.
36
Figure 28. Location of the Sierra Quemada Caldera and related features
(modified from Henry and Price, 1986).
37
Figure 29. View of the Sierra Quemada from the South Rim.
Rocks that were formed prior to caldera collapse include minor lava flows
in and near the caldera, tuffaceous material below the Mule Ear Spring Tuff, and
dikes found north and west of the caldera including those that can be seen along
Ross Maxwell Scenic Drive. The tuffaceous rocks can be at least as thick as the
Mule Ear Spring Tuff in places and are generally less-resistant, gently-sloping
outcrops above the Bee Mountain Basalt or below the Tule Mountain
Trachyandesite. Post-caldera activity includes the crudely-concentric ring
fracture and resurgent intrusions and related dikes as well as minor extrusive
materials. Dikes that extend northeast of the caldera appear to be related to this
stage of activity.
The most-widespread unit erupted from the Sierra Quemada Caldera, and
perhaps the most widespread volcanic unit in the park, is the Mule Ear Spring
Tuff which has been dated by others at about 34 Ma. It is thickest inside the
caldera where it exceeds 130 m (400 ft) and has clasts in excess of 10 m (33 ft).
Outside the caldera it thins away from the caldera margins from a maximum of
about 47 m (155 ft), through many areas with thicknesses around 25 m (80 ft), to
a more typical value of 6 m (20 ft). It can be seen in almost all of the areas
mapped as the Mule Ear Spring Tuff by Maxwell and others (1967), such as on
Goat and Kit Mountains, and also in a number of areas mapped as
“Undifferentiated Lavas” or other units. Most notably, it is thickest on the hill
about a mile northeast of Sotol Vista where it is present as three distinct layers
that are ash-flow tuffs, and on the smaller, rounded hill across Blue Creek
canyon about a half mile to the east-southeast.
38
Figure 30. This approximately 4.5 mi long total magnetic intensity profile starts
at A south of the caldera and runs about N – S up the Smoky Creek pack trail, to the
Dodson Trail, East along the Dodson Trail about one mile and then SE up hills within the
caldera. The increase in signal from A – B is the result of approaching the deep intrusive
masses of the caldera. Between B and C the trail is within a narrow canyon cut into the
ring-fracture-intrusion and the data show the short wavelength/high amplitude
characteristics of the caldera intrusives. The trail from C to D crosses the complexly
faulted (caldera collapse) lower portion of the interior of the caldera, which itself is down
faulted along a Basin and Range fault from the upper portion at D. At D the intensity
increases sharply as the trail crosses to the relatively uplifted upper portion of the caldera,
this increase is probably indicative of the decreased depth to the buried intrusive masses
of the upper portion of the caldera. In the stretch D – E the transect crosses a few dikes
and climbs to the intersection with the Dodson Trail (near E). The drop in average
intensity before E may be the result of the trail climbing away from the buried intrusives.
From E to F the path is mostly within caldera collapse debris and younger (reworked)
sediments. The relatively constant data are consistent with this geology with the
exception of two major excursions where dikes are crossed. We have more than 80 miles
of such profiles over and around the caldera.
Stop 7b. TERLINGUA UPLIFT OVERVIEW
Richard J. Erdlac, Jr., Department of Geosciences, University of Texas Permian
Basin, Odessa, Texas.
The Terlingua monocline forms the southern boundary of a much larger
structural feature called the Terlingua uplift (Figure 31). The uplift is roughly
rhombic in shape and is both structurally and topographically high (Erdlac, 1990).
This feature extends 29 km north-northwest and 23 km in maximum width. The
uplift is bounded to the west by the Fresno monocline that joins with the
Terlingua monocline at the southwest corner of the uplift (Figures 31 and 32).
The Fresno monocline strikes northwest into the southwestern edge of the
Solitario, which lies within the northwestern part of the uplift. It overprints and
obscures any northwesterly continuation of the Fresno monocline. Eastern and
northern boundaries of the uplift are more obscure. Cretaceous rocks dip gently
(less than 25o) northeast off the northeast flank and are cut by a series of
39
northwest-striking Basin and Range faults. The northern uplift boundary may
coincide with the Tascotal Mesa fault and is largely buried by Tertiary strata.
Figure 31. This Landsat image shows the Terlingua uplift (circled area) within the
context of numerous other features within the Big Bend region, both within Texas and
Mexico. For example the dark area to the west represents the Bofecillos Mountains. To
the south in Mexico lie the San Carlos and Sierra Rica caldera complex. The
southeastern edge of the image displays reddish rocks that comprise the Chisos
Mountains. Finally, the small dark circular areas with the central part of the image
represent numerous igneous intrusives found in the region.
Lower to Upper Cretaceous rocks crop out across the uplift. Santa Elena
Limestone is exposed over much of the uplift, whereas Del Rio Clay through
uppermost Cretaceous strata (Javelina Formation) crop out along the flanks or
are preserved locally in collapse structures or on the down-thrown side of faults
on the uplift.
The Terlingua uplift originated as a Laramide push-up structure (Erdlac, 1988,
1990; Erdlac and Erdlac, 1992) that predated the Tertiary intrusion and volcanic
activity that formed the Solitario. Physiographically, the uplift is divisible into a
southern half that displays numerous northeast- and northwest-striking faults and
a northern half upon which the Solitario is superposed. This superposition was
first noted by Baker (1934) when he stated that the Solitario rests on and merges
into a broad northwest-trending anticline, the Terlingua uplift. Field evidence for
Laramide deformation on the southern half of the uplift consists of folds, faults
(strike-slip, thrust, and contraction), tectonic stylolites, and various stratigraphic
relationships (Erdlac, 1990, 1994).
The faults that cut across the southern part of the Terlingua uplift are highly
orthogonal in nature (Figure 32). Many of these faults display both strike-slip as
40
well as dip-slip movement. The northeast-trending faults are interpreted as
Laramide in age whereas the northwest-trending faults are of Basin and Range
age. As the northeast-trending faults intersect the Terlingua monocline along the
southern edge of the uplift, the monocline displays a noticeable left-stepping en
echelon pattern in its strike (Figure 32).
Figure 32. This Landsat image shows the extent of the Terlingua uplift push-up
structure. It lies in the central and western (left) part of the image. This rhombic-shaped
feature is bounded on the south by the Terlingua uplift. The circular Solitario lies astride
the uplift along its northwestern portion. The Fresno monocline forms the western
boundary of the feature, and is itself overlain by the western edge of the Solitario. The
darker areas along the northeastern, southeastern, and western edge of the image
represent volcanics and intrusive rocks. The very northern part of Mesa de Anguila lies
along the southcentral edge of the image. The light colored areas are composed of
Boquillas and Pen Formations. The brownish to grayish-blue color of the Terlingua
uplift represents the Santa Elena Limestone.
Younger Cretaceous strata of the Boquillas Formation predominate
immediately south of the uplift. The area between the uplift and Mesa de Anguila
forms a noticeable structural and topographic saddle (Figures 32 and 33). This
saddle rises to the north and to the south but dips gently to the east and west.
The overall significance, if any, of this saddle is presently unknown.
41
Figure 33. From Sotol Vista looking west and northwest, the Rio Grande is seen
cutting through and separating Sierra Ponce to the left and Mesa de Anguila to the right
(north). Santa Elena Canyon is the light V-shaped notch at the left edge in the center of
the image. Mesa de Anguila trends northward where it dips into the subsurface. A low
region shaped as a saddle separates Mesa de Anguila from the Terlingua uplift farther
north. The uplift is along the horizon in the central part of this image.
39.4
Exit Sotol Vista; turn right back on Ross Maxwell Scenic Drive.
39.5 Blue Creek Ranch Overlook. This is also the western trailhead for the
Dodson Trail, which passes through the Sierra Quemada before ending at
Juniper Canyon on the other side of the park, and the Blue Creek Trail, which
goes up into the Chisos Mountains and joins the Laguna Meadow – South Rim
Trail.
End of Day 1. Continue driving north on Ross Maxwell Scenic Drive. Turn left
(west) on the main park road to return to Study Butte, then north to the Longhorn
Ranch Motel, where we will check in for the night and set up camp.
42
DAY TWO
Return to the western entrance of Big Bend National Park. Reset your odometer
to zero.
0.0 Big Bend National Park, western entrance.
9.2 Junction with Ross Maxwell Scenic Drive.
19.4
Chisos Basin Junction. Turn right (south).
24.5 Lost Mine Trailhead. The Lost Mine trail provides the best, shortest route
into the high Chisos Mountains. Additionally, this trail takes the hiker past Casa
Grande and up the eastern side of and into the Pine Canyon caldera.
25.4
Campground entrance. Continue straight ahead.
25.6
Chisos Basin parking lot. Park for Stop 8.
Stop 8. EXTRACALDERA VENTS OF THE PINE CANYON CALDERA
John C. White, Department of Geology, Baylor University, Waco, Texas.
The eastern side of the Chisos Basin consists of a ridge of the highest
peaks in the Big Bend, including (from north to south), Casa Grande, Toll
Mountain, Emory Peak, and Ward Mountain (Figures 34, 35, and 36). Ward
Mountain is an intrusion of quartz syenite and is not considered further in this
discussion. The other peaks, however, consist of accumulations of extrusive
rock (South Rim Formation) that have been variably interpreted:
Figure 34. Casa Grande volcanic dome.
43
Figure 35. Toll Mountain volcanic dome.
Figure 36. Emory Peak volcanic dome.
Maxwell and others (1976) interpreted these peaks as representing
successive flows of ash-flow tuff and quartz trachyte to rhyolite lavas that erupted
from the Pine Canyon volcanic center through a paleovalley that extended from
the high Chisos Mountains to the present-day Castolon Graben (Burro Mesa,
Goat Mountain, etc.).
Ogley (1976) retained Maxwell and others’ (1967) interpretation, excepting
for Casa Grande, which he interpreted as a rhyolite lava dome. In his mapping of
the Pine Canyon volcanic center, he was the first to recognize it as a caldera and
named it the “Pine Canyon caldera”.
Barker and others (1986) recognized that the units of the South Rim
Formation in the high Chisos Mountains erupted from a different area than the
44
units of the South Rim Formation in the Castalon Graben, and also recognized
that Emory Peak is most likely a separate vent area for the rhyolite that forms the
peak, but otherwise retained Ogley’s (1976) interpretation. This recognition of
the different character of the two different South Rim Formations led Barker and
others (1986) to (informally) revise the stratigraphic nomenclature of the South
Rim Formation in the high Chisos (Table 3). Barker and others also recognized
that Emory Peak represents a separate volcanic vent, and the “Burro Mesa”
rhyolite of Emory Peak (renamed the “Emory Peak Rhyolite” by Urbanczyk and
White, 2000) may be a rheomorphic tuff due to the apparently gradational contact
between the rhyolite’s pumiceous base and massive interior easily seen on the
Boot Canyon trail.
Table 3. Units of the South Rim Formation.
Maxwell and others (1967)
Barker and others (1986)
High Chisos Units
Castalon Graben Units
Burro Mesa Rhyolite
Burro Mesa Rhyolite*
Burro Mesa Rhyolite
Lost Mine Member
Lost Mine Member
Wasp Springs Member
Wasp Springs Flow-Breccia
Boot Rock Member
Brown Rhyolite
Pine Canyon Rhyolite
n/a
* “Emory Peak Rhyolite” of Urbanczyk and White (2000)
Urbanczyk and White (2000) retained Barker and others’ (1986)
interpretation of the Pine Canyon Rhyolite, but reinterpreted the Boot Rock
Member and Lost Mine Member as representing maar-type surge deposits (and
occasional pyroclastic flows) and lava domes, respectively, that erupted along
the ring of the Pine Canyon caldera (see Stop 9b for a summary). They
interpreted the Casa Grande lava dome as a similar vent, but one that erupted
outside of the caldera (an extra-caldera vent). They retained Barker and others’
(1986) interpretation of the Emory Peak vent and also suggested that earlier
eruptions from the Casa Grande volcanic center were responsible for the
pyroclastic and lava flows that make up Toll Mountain and the South Rim.
Continued reconnaissance geology in the high Chisos Mountains has
revealed that the “ash-flow tuffs” of the external facies of the Boot Rock Member
beneath Toll Mountain (“The Pinnacles”) are, in fact, also maar-type surge
deposits similar in character to those found within the caldera and at Burro Mesa,
etc., elsewhere in the park (Figure 37). This strongly suggests that there are at
least three extra-caldera vents related to the Pine Canyon caldera: Casa Grande,
Toll Mountain, and Emory Peak. Which of these vents (if any) represent the
source for the tuffs and lavas that form the South Rim remains unknown.
45
Figure 37. The Pinnacles beneath Toll Mountain, as viewed from the Pinnacles
trail. The Pinnacles are the eroded remnants of maar-type surge deposits very similar to
other such deposits mapped variably as the Boot Canyon Member or Wasp Springs
Member. Note the downward dip of the unit from right to left.
After Stop 8, we return the way we came to the Basin Junction, then turn right
(east) towards Panther Junction (park headquarters). At Panther Junciton, reset
your odometer to zero.
0.0
Panther Junction
5.3
Junction with the Glen Springs backcountry road.
6.6
The Sierra del Carmen are at 12:00.
14.9
Pine Canyon Vista on the lefthand (east) side of the road. Stop 9.
46
Stop 9. PINE CANYON OVERVIEW
Figure 38. View from Stop 9. Crown Mountain, Lost Mine Peak, and Pummel
Peak define the outline of Pine Canyon (PC) and the Pine Canyon caldera. Hayes Ridge
in the foreground is a quartz-phyric peralkalic rhyolite that is part of the ring dike that
surrounds the caldera. Hayes Ridge continues (out of view) to the west through Juniper
Canyon before “disappearing” beneath Emory Peak—the rhyolite of which is
geochemically identical to Hayes Ridge, strongly suggesting that the Emory Peak
Rhyolite and vent area may represent a extra-caldera, post-collapse eruption of the Pine
Canyon ring dike.
Stop 9a. PINE CANYON WATERSHED RESEARCH PROGRAM
Kevin Urbanczyk Department of Earth and Physical Sciences, Sul Ross State
University, Alpine, and John Zak, Department of Biology, Texas Tech University,
Lubbock
The Pine Canyon Watershed Program was initiated during the summer of
1995 and is part of a network of monitored watersheds in National Parks and
Equivalent Reserves funded by the Biological Resources Division of the USGS.
The Pine Canyon research program represents the only long-term monitoring
program in the Chihuahuan desert along the US-Mexican border that examines
both abiotic and biotic contributions to watershed dynamics. The Pine Canyon
Watershed covers approximately 7,800 ha and extends about 19 km in an
47
easterly direction from the central Chisos Mountains. Permanent monitoring
sites have been established in the five vegetation zones that exist along the
watershed (see Figure 39): 1) Lost Mine - High elevation forests dominated by
Pinyon Pine and Live Oak (6900 ft, “A” on map), 2) Oak-juniper association in
upper Pine Canyon (6000 ft), 3) Sotol-grassland (5020 ft, “B” on map), 4)
Creosotebush association at Chilicotal Springs (3000 ft, “C” on map) and Rice
Tanks (3322 ft, “E” on map), and 5) Lowland Chihuahuan Desert scrub (2610 ft,
“F” on map), Glenn Springs location. Long-term research objectives of the
Program include the determination of precipitation and other meteorological data,
surface and groundwater hydrology and chemistry, trends in microbial activity
and soil nitrogen dynamics, decomposition and fungal diversity, and plant
species variations occurring due to anthropogenic inputs.
Significant findings to date include the observation that atmospheric
pollution is affecting the desert ecosystem. These atmospheric inputs include
nitrogen and sulfate that cause decreases in bacterial functional diversity and soil
pH values. These decreases in soil pH are also linked to an overall increase in
acidity of local precipitation.
Stop 9b. PINE CANYON CALDERA
Kevin Urbanczyk and John C. White, Department of Earth and Physical
Sciences, Sul Ross State University, Alpine
The Pine Canyon caldera is an Oligocene (31.7 to 32.9 Ma) volcanic
center in the Chisos Mountains in Big Bend National Park, Texas. It is located
within the eastern alkalic belt of the Trans Pecos Magmatic Province. Barker and
others (1986) described the Pine Canyon caldera as a down-sag caldera that
formed as the ground gradually subsided during the eruption of the Pine Canyon
Rhyolite rather than collapsing and producing a topographic scarp around the
margin of the caldera. This eruption was followed by a second eruption of ashflow tuff (the Boot Rock Member) in addition to quartz trachyte and rhyolite lavas
(Lost Mine Member), which comprise the South Rim Formation. The Pine
Canyon rhyolite has previously been interpreted as the caldera forming
ignimbrite, an interpretation that we retain. The Boot Rock member has been
interpreted as ash-flow tuff that filled the caldera (caldera fill facies), and
breached the southwest wall and flowed toward the southwest (outflow facies).
The Lost Mine member has been interpreted as a sequence of lavas and ashflow tuffs that erupted from the caldera and flowed to the southwest, but is not
currently found within the caldera. We instead reinterpret the caldera fill facies of
the Boot Rock member to be maar-type surge deposits occasionally overlain by
block and ash-flow deposits produced by a semicircular series of postcaldera ring
vents located along the margin of the Pine Canyon caldera. In a similar fashion
as documented at Burro Mesa, Goat Mountain, and Cerro Castellan (the Burro
Mesa "group"), these surge deposits demonstrate concentric inward dips toward
lava domes. This is particularly well exhibited along the arcuate ridge starting
with Crown Mountain in the south, and continuing counter clockwise to Pummel,
George Wright, Panther and Lost Mine peaks. These deposits do not directly
48
Figure 39. Map of the Pine Canyon research area.
49
correlate with the Burro Mesa group of eruptions, but appear to have resulted
from similar eruption mechanisms.
15.5
Eastern end of the River Road.
15.9
Lower Tornillo Creek Bridge
16.6
Junction with the Hot Springs road.
17.7
Junction with the Old Ore road.
18.3
Tunnel. The tunnel cuts through Santa Elena Limestone.
18.4
Rio Grande Overlook
19.1
Junction with Boquillas Canyon road; Turn Left.
20.5
Road to Boquillas, Mexico, crossing.
21.8
Junction with Boquillas Overlook road; Turn right.
22.3
Boquillas Overlook. Stop 10.
Stop 10. GEOMORPHOLOGIC CHANGES IN THE RIO GRANDE / RIO
BRAVO CHANNEL AND FLOODPLAIN
Carol Purchase, National Park Service, Big Bend National Park, Texas.
The Rio Grande / Rio Bravo river channel and floodplain have changed
markedly over the past 100 years. Irrigation diversions began to reduce low
flows as early as the 1700's in New Mexico. Elephant Butte dam first started
reducing flood flows in 1916. Currently, this part of the Rio Grande is
hydrologically disconnected with the Rio Grande in El Paso and above. All floods
from the U.S. portion of the river have been virtually eliminated by the many
dams in New Mexico and the small channel through El Paso. The river channel
disappears below Fort Quitman (60 miles below El Paso) into a large Salt Cedar
forest.
The Rio Conchos, which joins the Rio Grande/ Rio Bravo at Presidio,
Texas, has always provided the largest floods through this part of the river,
during late summer and early fall. Due to dams in this watershed, floods have
been reduced by about 50%. With the reduction in flows, the channel has
narrowed as the river is unable to transport the sediment deposited by tributaries
(this is especially evident at the mouth of Terlingua Creek at Santa Elena
Canyon, Stop 3). Salt Cedar, an exotic tree, now lines much of the river,
replacing native cottonwoods and willows. This tree colonizes the terraces
formed from the larger floods and then stabilizes the floodplain, reducing the
channel capacity, which increases the flood height of subsequent floods. River
50
cane (Arundo donax), another exotic, has also increased over the past several
decades, further narrowing the channel. Note the comparison between the
historic picture (the handout) of the river and the scene below this overlook
(Figure 40). The broad sandy riverbed is gone, replaced by a vegetated canal.
Floods today would have a hard time scouring the riverbanks to restore the
natural river geomorphology.
Figure 40. View of the Rio Grande / Rio Bravo looking downstream from the
Boquillas Overlook.
22.8
Junction with main road; Turn right.
23.7
Boquillas Canyon parking lot. Stop 11.
Stop 11. RIO GRANDE HOLOCENE VALLEY FILL
Rolfe D. Mandel, Department of Geography, University of Kansas, Lawrence
Near its terminus at the entrance to the Boquillas Canyon parking area,
Boquillas Canyon Road winds along the margin of the valley floor of the Rio
Grande. The road crosses an arroyo before it turns into the parking lot at the
head of the Boquillas Canyon trail. This arroyo has deeply dissected the
Holocene valley fill of the Rio Grande on the south side of the road.
At Stop 11 (Figure 1), we will examine thick deposits of fine-grained
Holocene valley fill beneath the floodplain of the Rio Grande. In 1996, a single
cutbank was cleaned with a shovel and described (Mandel, 1996). Four buried
soils were identified in the upper 342 cm of the valley fill (Figure 41 and Table 4).
These buried soils have weakly developed A-Bw profiles and, with the exception
of a bed of gravel in the C2 horizon of the third buried soil, are composed of
calcareous, silty clay loam.
Humates from the upper 10 cm of the Ab3 horizon yielded a radiocarbon age of
3,140+60 yr B.P. (Tx-8308). This age suggests that most, if not all, of the Table 4.
51
Table 4. Description of the section at Stop 11.
______________________________________________________________________________
Landform: Valley floor of the Rio Grande
Slope: 1 percent
Drainage class: Well drained
Vegetation: Creosote bush, mesquite, prickly pear, bristlegrass, and fluff grass
Described by: Rolfe Mandel
Date described: September 28, 1995
Remarks: Humates from the upper 10 cm of the Ab3 horizon yielded a radiocarbon age of 3,140+60 yr B.P. (Tx-8308).
______________________________________________________________________________
Depth
Soil
(cm)
Horizon
Description
0-10
A
Pale brown (10YR 6/3) light silty clay loam, brown (10YR 5/3) moist; weak fine
subangular blocky structure parting to weak medium and coarse granular; slightly
hard, very friable; common fine roots; violent effervescence; clear smooth boundary.
10-58
C
Laminated pale brown (10YR 6/3) and brown (10YR 5/3) silt loam and light silty
clay loam, brown (10YR 5/3) and yellowish brown (10YR 5/4) moist; massive;
slightly hard, very friable; laminae are 1-3 mm thick; few fine and very fine roots;
violent effervescence; abrupt smooth boundary.
58-71
Ab1
Pale brown (10YR 6/3) light silty clay loam, yellowish brown (10YR 5/4) moist;
weak medium prismatic structure parting to weak fine subangular blocky; slightly
hard, friable; violent effervescence; gradual smooth boundary.
71-106
Bwb1
Pale brown (10YR 6/3) silty clay loam, brown (10YR 5/3) moist; very weak fine
subangular blocky structure; slightly hard, friable; violent effervescence; abrupt
smooth boundary.
106-123
Ab2
123-221
Bwb2
221-238
Ab3
Brown (10YR 5/3) silty clay loam, yellowish brown (10YR 5/4) moist; weak
medium prismatic structure parting to weak fine subangular blocky; slightly hard,
friable; violent effervescence; gradual smooth boundary.
238-260
Bwb3
Very pale brown (10YR 7/3) silty clay loam, pale brown (10YR 6/3) to yellowish
brown (10YR 5/4) moist; weak medium prismatic structure parting to weak fine
subangular blocky structure; slightly hard, friable; violent effervescence; abrupt
smooth boundary.
260-293
C1
Laminated pale brown (10YR 6/3) and brown (10YR 5/3) silty clay loam and silt
loam, brown (10YR 5/3) and yellowish brown (10YR 5/4) moist; massive; slightly
hard, very friable; violent effervescence; abrupt smooth boundary.
293-310
C2
Stratified coarse and fine gravel; single grain; loose; clasts are well rounded to
subrounded; abrupt smooth boundary.
310-328
Ab4
Pale brown (10YR 6/3) silt loam to light silty clay loam, yellowish brown (10YR
5/4) moist; weak medium prismatic structure parting to weak fine subangular blocky;
slightly hard, friable; violent effervescence; gradual smooth boundary.
328-342+
Bwb4
Very pale brown (10YR 7/3) silty clay loam, pale brown (10YR 6/3) to yellowish
brown (10YR 5/4) moist; weak coarse prismatic structure parting to weak medium
subangular blocky structure; slightly hard, friable; violent effervescence.
Brown (10YR 5/3) light silty clay loam, yellowish brown (10YR 5/4) moist; weak
medium prismatic structure parting to weak fine subangular blocky; slightly hard,
friable; violent effervescence; gradual smooth boundary.
Pale brown (10YR 6/3) silty clay loam, brown (10YR 5/3) moist; very weak fine
subangular blocky structure; slightly hard, friable; common lenses of fine gravel at a
depth of 214-221 cm; violent effervescence; abrupt smooth boundary.
52
Figure 41. Section at Stop 11.
53
alluvium in the cutbank was deposited during the late Holocene. The buried soils
indicate that late-Holocene alluviation was punctuated by episodes of floodplain
stability. However, weak soil development (thin A-Bw profiles) during this period
suggests that each episode of stability was relatively short, perhaps lasting a few
hundred years.
Although the stratigraphy of Holocene valley fill at Stop 11 is more
complex than the stratigraphy observed beneath distal segments of the floodplain
of the Rio Grande elsewhere on the east side of the park, the alluvial chronology
is fairly consistent from one locality to the next (Mandel, 1996). For example,
where San Vicente Arroyo has dissected the floodplain of the Rio Grande
adjacent to River Road, humates from a deeply buried soil yielded radiocarbon
ages of 3,430+70 yr B.P. (Tx-8310) and 3,470+50 yr B.P. (Tx-8309) (Mandel,
1996). These radiocarbon ages, combined with the radiocarbon age determined
on the soil at Stop 11, suggest that the floodplain of the Rio Grande was
relatively stable around 3,400-3,100 yr B.P. Renewed aggradation after ca. 3,100
yr B.P. resulted in deep burial of former stable surfaces. This is an important
aspect of late Quaternary landscape evolution, and, as will be noted at Stop 11,
is crucial to the interpretation of the archaeological record of the Big Bend region.
27.3
Junction with main road; Turn left towards Rio Grande Village.
Rio Grande Village Store. You may stop here for gasoline, supplies,
restroom, etc. The main caravan will turn right to the Rio Grande Village picnic
area for lunch.
28.3
29.1 Rio Grande Village Picnic area. LUNCH STOP. After lunch, we will
backtrack to the Rio Grande Village store and then return to Panther Junction.
29.8
Rio Grande Village store. Turn left.
49.9
Panther Junction. Turn right towards Marathon, Texas.
57.9
Fossil Bone Exhibit junction; turn right.
58.2
Fossil Bone Exhibit parking lot. Stop 12.
Stop 12. FOSSIL BONE EXHIBIT VISTA
Stop 12a. A VIEW OF THE TORNILLO GROUP
Thomas M. Lehman, Department of Geosciences, Texas Tech University,
Lubbock, TX
The Fossil Bone Exhibit sits atop a low northwest-trending sandstonecapped cuesta known as Exhibit Ridge. This homoclinal ridge interrupts what is
otherwise the featureless broad alluvial valley of Tornillo Creek, here about six
miles across, known as Tornillo Flat (the Spanish word "tornillo" refers to the
screw bean mesquite trees that grow here along the intermittent water courses).
54
Exposed along the valley walls, all around Tornillo Flat, are strata of Late
Cretaceous through Eocene age that Udden (1907) named the Tornillo "Clay".
These strata are composed primarily of recessively weathering multicolored
mudstone interbedded with conglomeratic sandstone lenses that hold up cuestas
and hogback ridges amidst low lying badlands. The mudstone intervals are rich
in swelling smectite clays subject to rapid erosion even in the arid climate of Big
Bend. For the most part, these easily erodible strata are buried under the
alluvium of Tornillo Creek, and only exposed along the sides of the valley
beneath the retreating erosional escarpments of the Rosillos, Sierra del Carmen,
and Chisos pediments. In the Tornillo Flat area, strata of the Tornillo Group dip
to the south-southwest except where they are bowed up along the edges of the
laccoliths that comprise McKinney Hills (to the east) and Grapevine Hills (to the
west), and where the section is repeated by a series of northwest-trending
normal faults.
The Tornillo Clay was studied in detail and mapped for the first time by
Maxwell and others (1967), who revised the name to Tornillo Group and
subdivided the strata into three formations; in ascending order these are the
Javelina Formation (named for Javelina Creek, which drains the northeast side of
Tornillo Flat), Black Peaks Formation (named for the Black Peaks - three small
intrusions on the east side of Tornillo Flat), and Hannold Hill Formation (named
for Hannold Hill, on the south side of Tornillo Flat where the park highway climbs
out of the valley). Although originally believed by Udden to be entirely Late
Cretaceous in age, the Tornillo Group was shown by Maxwell and others (1967)
to include strata of Paleocene and Eocene age as well. They attempted to place
the contacts between the three formations so that they coincided approximately
with the Cretaceous/Paleocene and Paleocene/Eocene boundaries. Their
geologic map of Big Bend National Park remains the most detailed depiction of
the distribution of these strata available today. As the formations were originally
defined, the Javelina Formation included strata primarily of Late Cretaceous age,
the Black Peaks of Paleocene age, and the Hannold Hill of early Eocene age.
The Fossil Bone Exhibit displays fossils found in the Hannold Hill Formation, near
the base of the formation, within the Exhibit Ridge Sandstone Member (the
resistant sandstone unit that holds up Exhibit Ridge). All of the Tornillo Group
sediments accumulated in fluvial channel, floodplain, and lacustrine
environments, and throughout deposition of the sequence, stream flow was to
the southeast (Lehman, 1985; Beatty, 1992; Hartnell, J.A., 1980; Rigsby, 1982,
1986).
In addition to the outcrops on Tornillo Flat, these strata are exposed in the
valleys of Dawson Creek and Rough Run Creek along the western side of the
Park, and in the valley of Juniper Draw and other tributaries of the Rio Grande
south of the Chisos Mountains. Although most of the exposures of the Tornillo
Group are found within Big Bend Park, small outcrops are also found in areas
north of the Park, west, and south of the Park in Mexico.
The Tornillo Group is of broad interest because it preserves the
southernmost succession of Late Cretaceous, Paleocene, and Eocene
continental faunas and floras known in North America, and these differ in many
55
respects from those found in northern latitudes (Schiebout, 1974; Schiebout and
others, 1987; Rapp and others, 1983; Standhardt, 1986; Runkel, 1988, 1990;
Lehman, 1987; 1997; Wheeler and Lehman, 2000; Lehman, 2001). For
example, the Javelina Formation preserves remains of a dinosaur fauna
dominated by the sauropod Alamosaurus and the giant pterodactyl
Quetzalcoatlus. Remains of these animals are not found in northern latitudes
(e.g., Wyoming, Montana, and Canada), where at the same time such animals as
the horned dinosaur Triceratops and the theropod Tyrannosaurus dominated the
fauna (Lehman, 2001). Such observations provide evidence for the existence of
latitudinal life zones during Late Cretaceous time. Similar differences are found
in the Paleocene and Eocene faunas of the Black Peaks and Hannold Hill
Formations. These strata also contain a record of the Cretaceous/Tertiary
boundary extinction event, the closest continental record of this event in proximity
to the presumed impact site at Chicxulub, Mexico (Lehman, 1989, 1990; Straight,
1996).
The Tornillo Group strata are also important in recording the progression
of the Laramide Orogeny in the southern Cordillera and development of an
intermontane basin, known as the Tornillo Basin (Wilson, 1970; Lehman, 1986;
Lehman, 1991). Detailed mapping of the Tornillo Group strata in areas around
Laramide structures, as well as sedimentary provenance and paleocurrent data,
indicate that compressional deformation began during deposition of the Javelina
Formation (70 to 65 Ma) to define the margins of the Tornillo Basin (Lehman,
1991). Strata within the Tornillo Basin were deformed during early Eocene time
(57 to 54 Ma) and again prior to middle Eocene time (54 to 51 Ma). There is little
evidence for compressional deformation after the middle Eocene (post 51 Ma).
Stop 12b. TORNILLO FLATS GRASSLANDS
Carol Purchase, National Park Service, Big Bend National Park, Texas.
The large flat valley along Tornillo Creek and the flats in the northern
portion of the park once supported lush stands of grass. Historical accounts
describe “endless fields of grass”, “stirrup high grass” and also mention settlers
used to harvest hay on these flats for their livestock. When this land became a
park, grass was scarce on these flats due to overgrazing during dry years (Figure
xx). The soil found here, Tornillo Loam, a silty clay loam is extremely fragile
once vegetative cover is lost.
58
Figure xx. Tornillo Flat, 1950.
In the past 50 years, many areas have recovered to creosote, and in some
areas, native grasslands have returned, however many areas have accelerated
erosion and are now on a trend towards severely eroded badlands (Figure xx).
The most severely eroded areas tend to be on areas with slightly higher slope or
associated with water diversions built by the ranchers to funnel water into stock
tanks. Many of these areas have now lost 2 to 16 inches of topsoil and have
probably lost the ability to recover to historical native grasslands. Over 1000
acres of gully systems have been mapped, and an estimated 10,000 more acres
are severely degraded in the park. These valleys were on a depositional trend
prior to the last century; gradually filling with sediment deposited by 9 Point Dray
and Tornillo Creek. Now the trend is reversed and the valleys are being eroded
away, on a trend towards badlands which will eventually extend across the valley
floor to the surrounding hills.
59
Figure xx. Typical eroded area with 2 to 12 inches of soil loss.
58.4
Junction with main highway. Turn right.
76.1 Persimmon Gap and Northern Entrance to Big Bend National Park. End
of Field Trip and Day 2. Thank you very much for attending and making this a
successful event! Have a safe trip back home.
60
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67