late cenozoic drainage development in the southeastern basin and

Lucas, S.G., Morgan, G.S. and Zeigler, K.E., eds., 2005, New Mexico’s Ice Ages, New Mexico Museum of Natural History and Science Bulletin No. 28.
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LATE CENOZOIC DRAINAGE DEVELOPMENT IN THE SOUTHEASTERN BASIN AND
RANGE OF NEW MEXICO, SOUTHEASTERNMOST ARIZONA, AND WESTERN TEXAS
SEAN D. CONNELL1, JOHN W. HAWLEY2 AND DAVID W. LOVE3
1
New Mexico Bureau of Geology and Mineral Resources, Albuquerque Office, New Mexico Institute of Mining and Technology, 2808 Central Ave. SE, Albuquerque,
New Mexico 87106 [email protected]; 2New Mexico Bureau of Geology and Mineral Resources, emeritus; 3New Mexico Bureau of Geology and Mineral
Resources, New Mexico Institute of Mining and Technology, 801 Leroy Place, Socorro, New Mexico 87801
Abstract–Comparisons of regional stratigraphic, sedimentologic, structural, and geomorphic data for Neogene
basins of the southeastern Basin and Range, particularly those basins connected by the Rio Grande and upper
Gila River, reveal downstream-directed drainage integration by basin filling and spillover across low-standing
topographic sills between adjacent basins. Stream capture from adjacent internally drained basins probably
played secondary, but locally important roles in the development of integrated drainages in the region. In the
northern part of the Rio Grande drainage basin, late Miocene streams from mountainous headwaters in the Rio
Chama and Sangre de Cristo Mountains formed the ancestral Rio Grande, which flowed into playa lakes at the
southern part of the Albuquerque basin. By early Pliocene time, the ancestral Rio Grande drained into southern
New Mexico, western Texas, and northern Mexico. During late Pliocene time, the ancestral Rio Grande flowed
across a low topographic sill and continued downstream into western Texas and northern Mexico. Similarly,
although less well documented, drainage associated with the Gila River integrated downstream across basin
divides and intervening ranges from New Mexico into southeastern Arizona in Plio-Pleistocene time. The progression of regional drainage integration for both river systems does not appear to coincide with major climatic
events, but might be associated with progressive filling of tectonically quiescent or slowly subsiding basins.
Climatic controls on fluvial discharge and deposition are reflected by increased caliber of Plio-Pleistocene
axial-river sediments. Regional stratigraphic correlations support a climatic link for river-valley incision. Incision of basin floors and the development of river valleys and inset fluvial terraces began between 1.2 Ma and
0.67Ma for the Rio Grande. The mechanisms of climatically induced incision are not clearly understood, but
probably relate to episodes of increased stream power that might be linked to the increased amplitude and
higher frequency of climatic changes that occurred during Pleistocene time.
INTRODUCTION
The southeastern Basin and Range province of New Mexico, southeastern Arizona, and western Texas contains a rich array of landforms
and deposits formed during the post-Miocene interval of diminishing
tectonic activity that was coupled with increasing magnitude (amplitude) of climatic oscillations that started during late Pliocene time (Hawley
et al., 1969, 1976, 2000, 2002; Menges and Pearthree, 1989; Morrison,
1991a; Chapin and Cather, 1994; Cather et al., 1994; Mack, 2004; Smith,
2004). This contribution summarizes numerous studies of landscape
development and depositional history of extensional basins of the southwestern United States during the past few million years. This overview
emphasizes the development and possible causes of integrated drainage
and eventual incision of basin fill and the formation of entrenched river
valleys and associated landforms of the Rio Grande and Gila River, the
two major fluvial systems in the southeastern Basin and Range (Fig. 1).
This contribution does not represent a comprehensive summary of all
basins of the southeastern Basin and Range, but rather is a review of
geomorphic and stratigraphic data for internally drained basins that have
been reasonably well studied and basins associated with the Rio Grande
and Gila River.
The Gila River and Rio Grande occupy somewhat different climatic regimes and have slightly different tectonic histories, so a comparison of how these river systems have evolved could reveal possible similarities and differences in mechanisms for alluvial-basin development in
semi-arid extensional settings. The upper Gila River (more so than the
Rio Grande) is not solely confined within alluvial basins, but locally
flows across ranges at structurally controlled drainage divides. After
examining the tectonic, climatic, geomorphic and geochronologic back-
ground related to these river basins, our summary continues with brief
overviews of internally drained basins of the late Miocene in order to
compare sedimentation patterns and rates for basins before they became
influenced by throughgoing drainage. After examining the timing of fluvial integration among basins of the Rio Grande and Gila River, we
examine the timing of drainage incision. We conclude with paleogeographic reconstructions of drainage integration and discuss the roles that
tectonics and climate might have played in the evolution of the presentday integrated drainage throughout the southeastern Basin and Range
and Rio Grande rift.
Regional Tectonic Setting
The southeastern Basin and Range Province of New Mexico and
Arizona and its narrow continuation northward into the Southern Rocky
Mountains Province is characterized by tectonically extended terrane
that was broken into elongate, tilted fault-block uplifts that became
separated by full-graben and half-graben basins. Widespread, but areally limited volcanism is common in most (but not all) of the basins.
Extension in this region began roughly 20-30 million years ago and
continues presently along normal faults in New Mexico, Arizona, Colorado, and northern Chihuahua and Sonora (Chapin and Cather, 1994;
Pearthree, 1998; Machette et al., 1998). A major part of the Basin and
Range Province in New Mexico is the south-trending Rio Grande rift, a
major continental extensional structure separating the Transition Zone,
Colorado Plateau and part of the Southern Rocky Mountains Province
to the west from the eastern edge of the Basin and Range, Great Plains,
and eastern prong of the Southern Rocky Mountains provinces. Most
extensional basins in the region are asymmetric and are dominated by
half-graben forms. Symmetrical (full) grabens are also present, but are
126
FIGURE 1. Shaded-relief map showing locations of the southeastern Basin and Range, Great Plains, and Southern Rocky Mountain (SRM) provinces, and major highways
and drainages. The southeastern Basin and Range is divided into the Mexican Highland, Transition Zone, and Sacramento sections (Hawley, 1986). The Rio Grande rift
structural province is denoted by the stippled pattern. Basins of the Rio Grande include the San Luis (SLB), Española (ESB), Albuquerque (ALB), Socorro (SB), Palomas
(PB), Jornada (JB), Mesilla (MB), Tularosa (TB), and Hueco Bolson (HB). Internally drained basins include the Plains of San Agustin (PSA), Eagle Flat (EF), Mimbres
(MIB), San Bernardino (S), Animas (A), and Playas (P), and Salt (SL) basins. Basins associated with the Gila River include the Mangas Trench (MT), Safford-San Simon
(SS), and San Pedro Valley (SP) basins. Ranges include the White (wm), Gila (gm), Burro (bm), Tres Hermanos (th), Hatchet (ha), Peloncillo (pe), Chiricahua (ch), and
Piñaleno (pi) Mountains.
less common than their asymmetrical counterparts. Isolated basins commonly develop during early stages of extension (Chapin and Cather,
1994). As rifting progressed and rift-flank uplifts grew in stature, basin-bounding faults overlaped and created zones of strain accommodation between adjacent basins (Rosendahl, 1987). These zones of strain
accommodation occur where half-graben basins alternate domains of
tilt polarity along the rift zone strike (Muehlberger, 1979; Rosendahl,
1987; Faulds and Varga, 1998). The relief of these accommodation zones
is largely controlled by the degree of structural overlap among basinbounding fault segments (Rosendahl, 1987; Faulds and Varga, 1998).
Structural linkage of these extensional basins tend to overlap and enhances the topographic and hydrologic linkage among basins that are
axially aligned (Frostick and Reid, 1989; Leeder and Jackson, 1993),
as in the case of the Rio Grande rift. This resultant morphology creates
topographically connected basins with low bedrock divides separating
adjacent basins. Decreased throw across normal faults near accommodation zones also minimizes the structural and topographic relief (Faulds
and Varga, 1998) and can allow for drainage integration as low-lying
divides are filled.
Basins of the Rio Grande rift, where they extend into the southern Rocky Mountains Province, are distinctive because long, high-relief ranges, such as the Sangre de Cristo Mountains (Fig. 1), bound
them and create few places for interbasinal drainage to occur. The southern Rio Grande rift has a less prominent topographic expression, but is
recognized by higher heat flow, thinner crust, and locally extensive
volcanism (Seager and Morgan, 1979; Keller et al., 1990; Chapin and
127
tively ice-free, conditions (Zachos et al., 2001). The Pleistocene Epoch
began at 1.8 Ma (Berggren et al., 1995) and is marked by the presence
of large polar ice sheets and frequent, large-amplitude, cold-warm oscillations. Widespread ice-sheet development in the northern hemisphere
and strongly oscillating climatic conditions are recognized from both
marine and terrestrial records as early as 2.7-2.5 Ma (e.g., Morrison,
1991b; Smiley et al., 1991; Zachos et al., 2001). Long-term continental
paleoclimate proxy records are relatively rare, but generally correspond
to climatic events recorded in deep-sea sediments (e.g., Smiley et al.,
1991; Thompson, 1991; Kashiwaya et al., 2001; Zachos et al., 2001).
By about 0.9-0.8 million years ago, climatic proxy records show changes
in amplitude and frequency attributed to intensification of glacial-interglacial cycles at roughly 100 kyr intervals (Morrison, 1991b). The
influence of these climatic events on the stratigraphic and paleoecologic
record is preserved in the Basin and Range province (cf. Morrison,
1991a; Thompson, 1991). Even allowing for significant amounts of
intraregional and local variability, the marked climatic oscillations of
interglacial-glacial cycles appear to have dominated the forcing of
surficial geomorphic processes for at least half of the Pleistocene
(Hawley et al., 1976; Gile et al., 1981; Morrison, 1985, 1991a; Imbrie
et al., 1993; Winograd et al. 1997).
Geomorphic and Depositional Settings of Extensional Basins
FIGURE 2. Summary of the late Neogene time scale, including part of the
geomagnetic polarity time scale (Cande and Kent, 1995), North American Land
Mammal “Ages” (including the Rancholabrean, RLB; Woodburne, 1987). Temporal
distributions of d18O of benthic foraminifera in deep-sea cores from the equatorial
east Pacific (ODP 846) and equatorial east Atlantic (ODP 659) oceans (Haug and
Tiedeman, 1998; Shackleton et al., 1995). Major shifts in climate are recognized at
approximately 2.5 Ma and 0.8 Ma (cf. Zachos et al., 2001) and are shown as
horizontal gray bands. During the early Pleistocene, the d18O curves indicate a change
from a 41-kyr frequency to 100 kyr that occured near the beginning of large
continental glaciations. Vertical dashed lines approximate late Holocene d18O values.
Paleoclimatic proxy data for the western United States (Thompson, 1991),
southeastern Arizona (Smith, 1994), the eastern United States (Groot, 1991), and
major continental glaciations (illustrated by narrow gray bands), illustrates interpreted
changes recorded in terrestrial environments.
Cather, 1994; Pearthree, 1998; Machette et al., 1998). Ranges of the
southern Rio Grande rift and adjacent Basin and Range also tend to be
shorter and lower in elevation than the mountain ranges of the northern
rift.
As presently understood, rates of extension across the southeastern Basin and Range Province and along the Rio Grande rift started
slowly, increased during mid-Miocene time, and subsequently slowed
during Pliocene and Pleistocene time (Chapin and Cather, 1994). Nonetheless, the presence of Pleistocene fault scarps indicates that extension continues across the whole province (Pearthree, 1998; Machette
et al., 1998), albeit at a relatively slow rate.
Regional Paleoclimatic Setting
The late Cenozoic marks the most recent period of a long-term
cooling trend that culminated with the development of extensive glaciation of the northern hemisphere during Pleistocene time (Fig. 2).
The Pliocene and Pleistocene episodes constitute times of high-amplitude oscillations in the global climate that gave rise to alternating episodes of widespread cooling and glaciation followed by warmer, rela-
Nonmarine basins provide a record of sedimentation from upland catchments into lowland basins. In sedimentary basins, the preservation of sedimentary successions depends largely on tectonic subsidence and sediment flux (Leeder and Gawthorpe, 1987). The magnitude of sediment discharge on time frames greater than 103 or 104 yr,
and its partition into transverse and axial depositional components (see
below) plays an important role in basin-fill architecture (Blair and
Bilodeau, 1988; Gawthorpe and Leeder, 2000). Two broad categories
of basin fill are recognized in nonmarine extensional basins: basin floors
and basin margins. The basin floor, which occupies the lowest part of
the basin and typically extends along the long-axis of the basin, provides the best sedimentary evidence for drainage integration. Basinfloor settings contain deposits associated with internal (endorheic) and
external (fluvially integrated) surface-drainage conditions. The character of the basin floor is important in determining whether a basin was
subjected to external surface drainage or was internally drained. Finegrained fluviolacustrine sediments are common in internally drained
basins, whereas in externally drained basins, coarser grained fluviatile
deposits are present on the basin floor and commonly represent perennial rivers that flow through the basin. These are referred to as axialfluvial deposits.
Basin-margin deposits commonly come from local (i.e.,
intrabasinal) tributary streams that are typically oriented orthogonally
to the long axis of the basin (Fig. 3a-b). These are called transverse
deposits. Piedmont deposits represent locally derived transverse sediments associated with tributary streams draining fault-bounded uplifts
along the basin margin. In most extensional basins, faulting provides
excellent exposure of transverse-tributary deposits, which comprise the
bulk of the basin fill and are differentiated into alluvial-fan, alluvialslope, or fluvial-fan deposits. These deposits, however, provide little
direct information for differentiating internal from external basin
paleodrainage conditions.
Differential uplift and basin subsidence supply the potential energy for the sedimentary system and provides first-order control over
the driving forces of fluvial and glacial erosion across a variety of temporal and spatial scales. Structural controls at the junctions between
laterally adjacent or longitudinally (axially) linked basins and ranges
exert a strong influence on the topography of drainage divides. Low
divide relief allows for spillover of water and sediment from one basin
to another and controls the development of drainage in extensional settings (Frostick and Reid, 1989; Leeder and Jackson, 1993).
The ratio of subsidence to sedimentation rate controls the man-
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FIGURE 3. Photographs of landforms in the Basin and Range and Rio Grande rift: western front of Sangre de Cristo Mountains and low-relief plains of the Alamosa
(sub)basin, northern San Luis basin (a); localized internal drainage along western front of Manzano Mountains, a fault-bounded uplift along the eastern part of the
Albuquerque basin (b); incised tributary drainages cut into Plio-Pleistocene basin fill, northern flank of the Sandia Mountains, Albuquerque basin (c); western front of
Sandia Mountains, Albuquerque basin (Rincon Ridge in foreground), showing Sandia Wash, an incised tributary to the Rio Grande (d: courtesy of Grant Meyer).
ner of sedimentation in basins (e.g., Leeder, 1997). If subsidence is rapid
(relative to sedimentation rate), basins are internally drained and
underfilled. Lithofacies associated with underfilled basins are commonly fine-grained and are interpreted as relatively widespread
fluviolacustrine, playa-lake, or alluvial flat depositional environments.
Where sediment flux exceeds storage capacity, individual basins fill
and bury low-lying topographic divides, resulting in spillover into topographically lower adjacent basins (Gawthorpe and Leeder, 2000).
Fluvial deposition dominates basins where surface flows between basins are fluvially integrated (connected). If subsidence of the basin floor
is asymmetric, rivers tend to focus on the subsiding part of the basin,
resulting in the development of single axial river with asymmetric alluvial aprons along the margins of the basin (Mack and James, 1993).
Fluvial integration of adjacent basins can form by at least two
different mechanisms. Headward erosion of streams across a basin divide can capture drainage and integrate two formerly separate drainage
basins. Alternatively, overfilling of a basin can result in spillover into
an adjacent (lower) basin. Headward erosion can be problematic where
headwater areas are relatively small, although groundwater sapping
(caused by subsurface drainage between basins) could enhance headward
erosion. If headward erosion controls integration of formerly separate
basins, then the onset of through-going fluvial sedimentation should
become progressively younger upstream, or be more or less random. If
spillover is the dominant process, then the onset of through-going drainage should progressively decrease downstream.
During times of diminished (but not necessarily inactive) extension, sedimentation and expansion of drainage basins might eventually
provide enough sediment to aggrade basins to a level where they can
overtop low-relief topographic divides and connect to lower adjacent
lying basins. This results in the development of overfilled basins that
contain axially oriented drainage that convey extrabasinally derived
sediment through the basin. Externally drained (fluvially integrated)
basins are typically coarser-grained than their internally drained counterparts. This may be the result of open-system conditions brought about
by the transport of fine-grained sediments out of the basin as suspended
load. In contrast, internally drained basins trap all sediment.
Basin filling ceases when drainages incise and cut river valleys
(Fig. 3c-d). Where smaller drainages empty onto relict basin floor surfaces and have not yet integrated with the axial drainage (Fig. 3b), determination of entrenchment and the end of regional basin filling can be
difficult to distinguish (Connell et al., 2000, 2001a). Entrenchment leaves
a distinctive record of landforms and deposits as rivers episodically
incise into older fill (Fig. 3). In poorly drained areas, alluvial fans, associated with steep faulted mountain fronts, merge onto broad, low-relief
surfaces (Fig. 3b). In areas where drainages have not become integrated
with incised trunk rivers, streams debouche onto broad, low-relief surfaces that are not graded to the axial river.
Geochronology
Radioisotopic and biostratigraphic data greatly improve the chronological resolution of basin-fill. Radioisotopic ages of volcanic rocks
that are interbedded with the basin fill include mafic lava flows, ashflow tuffs, fallout ashes, and fluvially recycled volcanic clasts. Potassium-argon (K-Ar) and argon/argon (40Ar/39Ar) ages are reported from
numerous published and unpublished sources cited below and in the
Appendix. Earlier regional compilations of ages for the central and
southern parts of the Rio Grande rift have been completed by Bachman
and Mehnert (1978) and Seager et al. (1984). More recent work includes studies of radioisotopic studies of volcanic rocks interbedded
within the basin fill (e.g., WoldeGabriel et al., 2001; Smith et al., 2001;
Mack et al., 1998; Baldridge, 2004; Connell, 2004; Mack, 2004; Smith,
2004; Goff and Gardner, 2004; Wilks and Chapin, 1997; NMGRL,
1998). Vertebrate fossils have been collected from numerous sites and
are summarized in Morgan and Lucas (2003, this volume). Age assignments for vertebrate fossils are based on correlation to the provisional
North American Land Mammal “Ages” (NALMA) as discussed in
Woodburne (1987, 2004). Paleomagnetic studies also provide much
needed correlative age control for parts of the basin-fill succession (e.g.,
MacFadden, 1977; Lindsay et al., 1990a; Smith, 1994; Tedford and
Barghoorn, 1999; Mack et al., 1993, 1998; Geissman, 2004). Such studies use the ability of certain sediments to act as high-fidelity recorders
of variations of the earth’s magnetic field through time. Paleomagnetic
studies provide high-resolution temporal constraints on sedimentation
rates and allow correlations to be made of sediments within and among
basins.
A number of volcanic ashes have a wide distribution and provide excellent geochronologic markers (Luedke and Smith, 1978, 1991;
Sarna-Wojciki and Davis, 1991). Ages for some of these ashes have
been recently refined. Recent 40Ar/39Ar dates for tephra from Yellowstone
National Park are 0.64 Ma for the Lava Creek B ash, and 2.06 Ma for
the Huckleberry Ridge ash (Lanphere et al., 2002; cf. Izett et al., 1992).
The distributions of these two stratigraphically important tephra are
illustrated in Izett and Wilcox (1982) and Sarna-Wojciki and Davis
(1991). The Bishop ash, a widespread fallout tephra from eastern California, is present in New Mexico and has been recently dated at 0.76
Ma using the 40Ar/39Ar method (Sarna-Wojciki et al., 2000).
Other geochronologically useful tephra include eruptive products from the Jemez Mountains of north-central New Mexico. Two units
formed during the creation of two calderas in the Jemez Mountains; the
lower and upper Bandelier Tuffs, which have been dated at 1.61 and
1.22 Ma, respectively (Izett and Obradovich, 1994). Intra-caldera
eruptives of the Cerro Toledo Rhyolite have also been dated (Izett et
al., 1981) and provide locally important stratigraphic markers in the
Rio Grande rift (e.g., Mack et al., 1996; Connell, 2004, Smith, 2004).
Plio-Pleistocene volcanic rocks along the southeastern flank of the Jemez
volcanic field have also been dated (WoldeGabriel et al., 1996, 2001).
Regional Geologic Setting
The southeastern Basin and Range Province of New Mexico,
western (Trans-Pecos) Texas, southeastern Arizona, and northern
Mexico covered in this overview has a complicated geologic setting
and history. Hawley (this volume) describes the regional geomorphic
setting of New Mexico. The southeastern Basin and Range Province
contains somewhat evenly spaced, elongate, subparallel mountain ranges
and intervening alluvial basins that formed during late Cenozoic extension (Fig. 1; Dohrenwend, 1987; Morrison, 1991b). The southeastern Basin and Range has two physiographic subdivisions: the Mexican
Highland and Sacramento sections (Hawley, 1986). The Sacramento
section forms the broad physiographic and structural boundary zone
between the Rio Grande rift and the Pecos Valley section of the southern Great Plains Province (Fig. 1). The Mexican Highland section contains elongate ranges with broad, elongate basins that extend from northcentral New Mexico to southeastern Arizona. Basin-floor elevations in
the Mexican Highland section are between 1000 and 2100 m and most
mountains and summits of plateaus are between 2000 to 3300 m. The
Sonoran Desert section of the Basin and Range Province lies west of
the Mexican Highland section and is not described here.
Intermontane basin fills of the Mexican Highland section contain thick alluvial and lacustrine deposits. These basins occupy 60 to
80 percent of this section, and the intervening complex range blocks
include a wide variety of rock units of Proterozoic through Cenozoic
129
age. Alluvial, eolian, and lacustrine sediments, and a variety of interbedded
volcanic flows and tephra, form the primary depositional record for the
region.
Internal drainage characterizes intermontane basins of the Mexican Highland section between the Rio Grande rift and the Peloncillo
range along the New Mexico-Arizona border. Only the western and
northern parts of the Transition Zone and Mexican Highland have deep
valleys and well-integrated drainage systems. Widespread dissection
of basin fill has only occurred in areas where drainage has been integrated with regional (e.g., Rio Grande and Gila River) or local (e.g.
Mimbres and Animas) fluvial systems during latest Pliocene and early
Pleistocene times. Aggradation of central basin floors has continued in
the extensive (topographically) closed-basin systems of the Mexican
Highland (e.g., Animas-Lordsburg, Jornada del Muerto, Mimbres, Playas, and Tularosa). Many contain large pluvial lakes of Pleistocene
age, such as the Animas, Mimbres, Playas, and Hatchita-Moscos basins (Hawley, 1993; Kennedy et al., 2000; Gile, 2002; Allen, this volume).
The Transition Zone is a physiographic and structural province
that separates the Colorado Plateau from the Basin and Range region
to the south. This province has its own unique geologic history, having
undergone pre-Neogene uplift and basin formation, extensive volcanism, and some Basin-and-Range-style extension. It includes the upper
Gila River and the San Francisco River, a major headwater-tributary to
the Gila River.
The Rio Grande rift structural province lies within the eastern
Mexican Highland section of the Basin and Range (Figs. 1 and 4;
Hawley, 1978; Chapin and Cather, 1994) and extends into the Southern Rocky Mountains to the north. The Rio Grande rift is a chain of
axially (i.e., longitudinally) connected half-graben extensional basins
with opposing tilt polarity (Chapin and Cather, 1994). The rift sits between the Colorado Plateau to the west and Great Plains structural
provinces to the east and extends northward through south-central Colorado, where it is bounded by ranges of the Southern Rocky Mountains,
and southward, into the highlands of Chihuahua and Trans-Pecos Texas
(Woodward et al., 1978; Tweto, 1979; Chapin and Cather, 1994). The
Rio Grande is the axial river that follows this series of tectonically
active, longitudinally aligned basins that form its namesake rift (Fig.
4).
The headwaters of the Rio Grande are in the southern San Juan
and Sangre de Cristo Mountains (Fig. 4), which have experienced extensive and repeated glaciations during Pleistocene time. Much of the
moisture that drives this river is derived from winter precipitation,
mostly from snow pack in the headwaters region (Douglas et al., 1993).
The Gila River drains relatively high elevation regions of the White
Mountains of Arizona and the Mogollon Mountains of southwestern
New Mexico. Headwaters of the Gila River are slightly lower than
headwaters of the Rio Grande and have experienced only minor glacial
activity (Blagbrough, 1968, 1986, 1994; Péwé et al., 1984; Merrill and
Péwé, 1977). Both summer rain and winter snow provide moisture for
the upper Gila River region (Douglas et al., 1993).
The Rio Grande rift contains a much larger number of Plio-Pleistocene active faults than in the adjacent Mexican Highland section
(Machette et al., 1998; Pearthree, 1998). Slip-rates in the region are
generally less than about 100 m/Myr (0.1 mm/yr) during the past few
million years (Colman et al., 1985; Menges and Pearthree, 1989;
Machette et al., 1998; Koning and Pazzaglia, 2002), but were probably
much higher during the Miocene (e.g., Menges and Pearthree, 1989;
Chapin and Cather, 1994; Mack et al., 1994a). Geodetic surveys demonstrate continued crustal deformation in region (e.g., Larsen et al.,
1986; Fialko and Simmons, 2001).
BASINS OF THE RIO GRANDE RIFT
The Rio Grande is about 3000 km long and drains more than
675,000 km2 of the southwestern United States and Mexico. The Rio
130
FIGURE 4. Shaded-relief map illustrating the extent of Rio Grande drainage and
internally drained basins. The Continental Divide marks the western boundary of
this figure. Areas of through-going drainage associated with the ancestral Rio Grande,
during Pliocene and early Pleistocene time, are depicted by stippled pattern. The
Rio Grande drainage system is subdivided into northern contributory, central trunk
and distributary, and southern distributary sections. Basins of the Rio Grande include
San Luis, Española (ESB), Albuquerque (ALB), Socorro (SB), San Marcial (SMB,
Milligan Gulch basin of Hawley, this volume), Jornada (JB), Engle (ENB), Palomas
(PB), Coralitos-Hatch-Rincon (CR), Mesilla (MB), Tularosa (TB), and Hueco
Bolson (HB). Internally drained basins include the Plains of San Agustin (PSA),
Mimbres (MIB), and Bolson de los Muertos (LDMB). The San Luis Hills (slh)
separates the Alamosa (sub)basin (AL) from the Taos Plateau volcanic field (TPvf).
Ranges include the Sierra de las Uvas (uv), Black Range, Organ Mountains (om),
Sacramento Mountains, Ladron Mountains (lm), Florida Mountains (fl), Franklin
Mountains (fm), East Potrillo Mountains (ep), and San Andres Mountains.
Geomorphic features, such as the Sunshine Valley (sv), Llano de Albuquerque (lda),
La Mesa (lm), and Cambray fan (cf) are shown for reference. Drainages (shown by
white lines) are calculated using DEM and are only shown for reference.
Grande rift can be traced from central Colorado through New Mexico and
to West Texas and Mexico as a series of linked, en echelon basins, many
of which are alternating half-grabens with adjacent rift-flank uplifts.
From north to south, the major basins are the San Luis basin, Española
basin, Albuquerque basin, Socorro basin, San Marcial, Engle, and Palomas
basins, Hatch, Rincon, Mesilla, Jornada, and Hueco basins (Fig. 4).
Deposition within the Rio Grande rift began during late Oligocene
time (~25-7 Ma) within internally drained basins (bolsons) where streams
derived from emerging basin-margin uplifts terminated onto broad alluvial plains or ephemeral to intermittent playa lakes and alluvial flats
(Chapin and Cather, 1994). These bains do not contain extensive evaporite deposits, suggesting that many of them might be (partly) connected
by subsurface (groundwater) flow. As these internally drained basins
filled topographic divides were buried by late Miocene or Pliocene time
and the axial river (western Rio Grande) began to flow into southern New
Mexico (Fig. 5; Gile et al., 1981; Chapin and Cather, 1994; Mack et al.,
1993, 1998).
Drainage of the Rio Grande can be considered in terms of the
contributive and distributive drainage nets of Allen (1965; Lozinsky
and Hawley, 1991), where drainage is collected through a contributive
network of tributary streams, transferred through a trunk river, and
eventually emptied across a distributive drainage network. Headwater
basins contain tributaries that form the up-stream contributory (or
contributary) section. The San Luis, Española, and Albuquerque basins represent the northern contributory section, defined by the presence of rather large tributary drainages that feed into the main-stem
ancestral Rio Grande. A relatively short trunk section is present where
drainage is confined within narrow and elongate half-graben basins
containing few large tributaries. The Socorro, San Marcial, Engle, and
Palomas basins generally represent the central trunk-river section and
contain deposits typical of half-graben basins; however, this trunk-river
distinction is not clear everywhere. The southern distributary system is
recognized by repeated occupation of adjacent basins by the axial river
across relatively low-relief topographic (and structural) divides. Below
the Rincon area, (between Truth or Consequences and Las Cruces, New
Mexico), the river forms a quasi-distributary drainage pattern that episodically spills laterally into adjacent basins (Mack et al., 1997).
Entrenchment of the basin fill is obvious along the valley of the
Rio Grande. Former positions of the ancestral Rio Grande are recognized by the presence of inset terraces along the valley margins. These
terraces record when the axial river shifted from net incision to episodes of aggradation. Few of these terrace deposits have been dated;
however, many inset deposits have distinct geomorphic and pedogenic
characteristics that allow limited correlation along the Rio Grande system. The literature is too vast to include all such geomorphic and PlioPleistocene stratigraphic studies; however, the reader is referred to the
following for summaries in New Mexico: Gile et al. (1981), Machette
(1985), Dethier et al. (1988), Dethier and McCoy (1993), Pazzaglia
and Wells (1990), Connell (1996), Connell and Love (2001), and
Pazzaglia and Hawley (2004).
Santa Fe Group
The depositional fill of basins of the Rio Grande is the Santa Fe
Group, which was first named for deposits described near Santa Fe,
New Mexico (Hayden, 1869). Most workers recognized that, in addition to the physiographic and structural continuity of the Rio Grande
Valley, major lithologic components of basin fills are repeated among
basins of the Rio Grande rift (Bryan, 1938, Spiegel and Baldwin, 1963;
Hawley et al., 1969; Seager et al., 1971; Hawley, 1978; Chapin and
Cather, 1994). This consistency is particularly striking in Pliocene and
Pleistocene sections of the depositional record. The age of the Santa Fe
Group generally ranges from about 25 Ma to less than 1 Ma (Chapin
and Cather, 1994). Basin floors contain locally thick eolian, playa-lake,
fluviolacustrine, and fluvial sediments. As presently treated by most
workers in New Mexico (e.g., Hawley, 1978; Chapin and Cather, 1994;
Connell, 2004; Mack, 2004; Smith, 2004), the Santa Fe Group is a
succession of piedmont-slope and basin-floor sedimentary facies that
include debris-flow, alluvial-fan, alluvial-flat, playa-lake, fluvial-plain,
eolian, lacustrine, and paludal deposits. The Santa Fe Group also locally includes interbedded basaltic to silicic volcanic flows.
The termination of widespread basin filling was marked by the
beginning of long-term incision that ultimately led to development of
entrenched valleys and dissection of about 100-215 m of the basin fill
131
FIGURE 5. Correlation chart showing major stratigraphic units within selected basins of the Rio Grande rift and eastern Mexican Highland section of the southern Basin
and Range. Explanation of symbols is on Figure 9. Black rectangles denote volcanic rocks. Black triangles denoted dated tephra. Black ellipses indicate maximum ages of
fluvially recycled pumice gravel. Inset units (named in italics) are shown as terrace deposits (Qt, Qag), valley alluvium (Qa, Qr), or have formal lithostratigraphic names.
Data sources listed in appendix. Other units include the Totavi lentil (TL) of the Puyé Formation of Griggs (1964) and the gravel of Lookout Park (QTlgp) of Smith and
Kuhle (1998).
(Machette, 1985). The incised valley of the Rio Grande developed during alternating episodes of entrenchment and partial aggradation of the
river channel, resulting in the deposition of a suite of inset fluvial terrace fills (Fig. 5; Gile et al., 1981; Reneau and Dethier, 1996; Dethier,
1999; Smith et al., 2001; Connell and Love, 2001). Entrenchment of
the Rio Grande and major tributaries was well underway by middle
Pleistocene time (Dethier, 2001); however, aggradation locally continues along non-incised tributary drainages that have not become integrated with the Rio Grande (e.g., Gile et al., 1981; Pazzaglia and Wells,
1990; Connell et al., 2000). The timing of incision is well dated to
about 0.7-0.8 Ma in southern New Mexico (Mack et al., 1998). Elsewhere in the Rio Grande rift, incision is less well constrained, but is
considered to have begun before middle Pleistocene time (Connell, 2004;
Smith, 2004; Mack, 2004).
Northern Contributory Section
The northern contributory (contributary) section is characterized
by relatively large tributaries to the Rio Grande. This section spans the
headwater reaches of the Rio Grande, from the San Juan Mountains
and San Luis basin, southward through the Albuquerque basin (Fig.
5A-C). Rather large, high-elevation watersheds of the southern Sangre
de Cristo, Tusas, and Jemez Mountains provide discharge to the Rio
Grande headwaters region. Other large tributaries include the Rio Puerco
and Rio Salado, which enter the Rio Grande from the west near the
southern end of this section. The Rio Grande and major tributaries,
such as the Rio Chama, contain abundant rounded orthoquartzite gravels that are derived from northern New Mexico and southern Colorado.
San Luis Basin
The San Luis Valley (Upson, 1939; Lambert, 1966) is the surface
expression of the San Luis basin, the largest structural basin of the Rio
Grande rift (Chapin, 1971, 1987; Hawley, 1978; Lipman and Mehnert,
1979; Tweto, 1979). The basin extends about 220 km south, from
Poncha Pass in south-central Colorado to the Picuris Mountains and
Embudo fault zone in north-central New Mexico (Dungan et al., 1984). It
is flanked to the east by the Sangre de Cristo Mountains and to the west
by the San Juan, and Tusas Mountains. The basin is primarily an easttilted half graben, bounded by frontal faults of the Sangre de Cristo
Mountains, with local intrabasinal horst blocks (Brister and Gries, 1994).
The San Luis Hills, near the Colorado-New Mexico border separate this structural basin into two physiographic and volcano-tectonic
subdivisions (Keller et al., 1984; Kluth and Shaftenaar, 1994). The
mostly undissected and poorly drained Alamosa (sub)basin (Upson,
1939) lies north of the San Luis Hills and includes a large fluvial fan of the
upper Rio Grande, an eolian dune field (Great Sand Dunes National
Monument), and piedmont-slope surfaces that grade to an extensive
basin-floor alluvial plain dotted with marshes, ponds, and shallow alkali
lakes (Siebenthal, 1910). The Taos Plateau lies south of the San Luis
Hills and forms the southern part of the San Luis basin. The Taos Plateau
consists of a large volcanic field interbedded with basin-fill sediments
and cut by the Rio Grande Gorge.
The Alamosa (sub)basin contains the Plio-Pleistocene Alamosa
Formation, a mostly buried alluvial, fluvial and lacustrine succession that
is as much as 500 m thick (Fig. 5A; Powell, 1958; Burroughs, 1981;
Rogers et al., 1992). Paleoenvironmental, biostratigraphic,
132
FIGURE 6. Longitudinal profile of the Rio Grande (a; from Belcher, 1975), illustrating locations of towns and major basins along its course to the Gulf of Mexico (b).
Basins include the San Luis (SLB), Española (ESB), Albuquerque (ALB), Mesilla (MB), and Hueco (HB). The upper Rio Grande knickzone forms a prominent convexity
in the profile in the southern San Luis basin. This knickzone was formed by emplacement of voluminous basaltic rocks of the Taos Plateau volcanic field. The ancestral Rio
Grande began to cut into these basalts resulting in a decrease in incision upstream from nearly 200 m near Pilar, New Mexico, to less than 3 m at Alamosa, Colorado (c; after
Wells et al., 1987).
magnetostratigraphic, and tephrachronologic investigations of the upper
Alamosa Formation at Hansen Bluff provide the best-documented and
most comprehensive description in the region (Rogers et al., 1992). These
deposits include the 0.76 Ma Bishop ash. Lacustrine deposits and geomorphic features in the San Luis Hills supports the presence of a lake
that probably drained southward by middle Pleistocene time (Machette,
2003). Most tributary streams are poorly integrated with the Rio Grande
and form a weakly dissected landscape that is dominated by aggradation
(Pazzaglia and Wells, 1990).
The Sunshine Valley area of the largely undissected Costilla
Plains of Upson (1939) is east of the Taos Plateau and upper Rio Grande
Gorge and north of the Red River Canyon between Questa and the
Colorado Stateline. The tectonically active front of the central Sangre
de Cristo Mountains rises abruptly to the east (Menges, 1990) and the
Guadalupe Mountain-Cerro Chiflo uplift bounds Sunshine Valley on
the south. Near the northern part of the Taos Plateau, Winograd (1959,
1985) noted lacustrine beds resting on basalt that interfinger with alluvium derived from the Sangre de Cristo Mountains. These fine-grained
sediments probably were probably deposited in depressions formed by
the combined action of tectonic deformation and damming by flows of
the Taos Plateau volcanic field.
South of the San Luis Hills, the Taos Plateau is underlain by a
thick sequence of lava flows of the Pliocene Servilleta Basalt and
interbedded alluvial, paludal, and lacustrine deposits of the upper Santa
Fe Group. Flows and vents of the Taos Plateau volcanic field were
emplaced between 5 and 1 Ma and cover much of the southern part of
the San Luis basin (Lipman and Mehnert, 1979; Dungan et al., 1984,
1989; Appelt, 1998). Volcanic ash beds exposed in basin fill in basinmargin deposits along the eastern edge of the Taos Plateau are derived
from upper Bandelier and Cerro Toledo eruptions of the Jemez volcanic
field about 1.22 to 1.47 Ma (Izett, 1981; Izett et al., 1981; Heiken et al.,
1986; Izett and Obradovich, 1994).
South of the Sunshine Valley area, Plio-Pleistocene fluvial deposits lie within incised valleys associated with the Rio Grande gorge.
The constricted part of the Rio Grande gorge between Cerro Chiflo and
Guadalupe Mountain separates a relatively shallow Rio Grande channel (100 m) from the 200 to 300-m-deep lower canyon segment that
cuts across the southeastern Taos Plateau. A knickzone with over 245
m of relief is present along Rio Grande between Cerro Chiflo and
Guadalupe Mountain separates the relatively shallow valley of the Rio
Grande from a deep lower canyon segment that cuts across the southeastern Taos Plateau (Fig. 6). This resistant knickzone is interpreted to
have formed during emplacement of the Taos volcanic field. Integration of the Rio Grande headwaters did not occur until this knickzone
migrated north to drain the Sunshine Valley and Alamosa (sub)basin
areas (Wells et al., 1987).
River gravels cap high-level erosion surfaces along the Rio
Grande gorge are present as far north as the knickzone between the
mouth of Red River Canyon and the Guadalupe-Chiflo constriction
(Wells et al., 1987; Pazzaglia and Wells, 1990). Rather than extending
farther up the San Luis Valley these fluvial gravels swing to the east as
rim-capping deposits flanking the lower Red River Canyon. Gravel
composition and size indicates that these gorge-capping river gravels
clearly have a Red River provenance (Wells et al., 1987). River gravels in the Taos Plateau north of the Guadalupe-Chiflo constriction are
confined to intra-gorge surfaces. Incision of the southern San Luis basin occurred as the Rio Grande cut into basaltic rocks of the Taos volcanic field. By middle Pleistocene time, much of the gorge had already been
cut (Fig. 6; Wells et al., 1987).
Española Basin
The Española basin is a west-tilted half graben bounded on the
east by the faulted dip slope of the southwestern Sangre de Cristo Mountains and on the west by the Pajarito fault zone and other geologic
structures now obscured by the Jemez volcanic field (Smith et al., 1970;
Bailey and Smith, 1978; Kelley, 1978; Manley, 1979a, b, 1984;
Golombek, 1981; Gardner et al., 1986; Smith, 2004). The Española
Valley is largely an erosional feature that was cut deeply into deformed
strata of the Santa Fe Group by drainages of the Rio Grande, Rio Chama
(Reneau and Dethier, 1996), and tributaries from the Sangre de Cristo
Mountains. In the Abiquiu embayment, just north of the Jemez Mountains, the Española basin extends across the Embudo-Pajarito fault zone
to Neogene faults bounding the Colorado Plateau and Southern Rocky
Mountains uplift (Tusas Mountains) in the Abiquiu-Cañones area north
of the Jemez Mountains (Dethier and Manley, 1985; Aldrich, 1986;
Smith et al., 2002). The basin ends to the southeast in the Santa Fe
embayment near Santa Fe, New Mexico (Kelley, 1978).
The Plio-Pleistocene depositional record of the Española basin
has been well documented (Fig. 5B; Spiegel and Baldwin, 1963; Galusha
and Blick, 1971; Hawley, 1978; Baldridge et al., 1984; Gardner et al.,
1986; Heiken et al., 1986; Ingersoll et al., 1990; Waresback and
Turbeville, 1990; Self et al., 1996; WoldeGabriel et al., 2001; Koning
et al., 2002). The Jemez volcanic field (active since middle-Miocene
time) obscures important stratigraphic and structural relationships;
however, some inferences can be made regarding the drainage history
of this basin. There appears to be little evidence for extensive
fluviolacustrine sedimentation in early Miocene time (Boyer, 1959).
The presence of coarse-grained axial river deposits is well represented
by the Plio-Pleistocene Totavi lentil of the Puyé Formation, which contains quartzite-bearing gravel that records the presence of the ancestral
Rio Grande in the Española basin (Griggs, 1964; Manley, 1979a;
Waresback and Turbeville, 1990). Manley (1979a) suggested that axial
drainage of the ancestral Rio Grande began after deposition of the Cejita
Member (upper Miocene) of the Tesuque Formation, which is also dominated by fluvial sediments (Koning, 2003). Water-supply wells drilled
in the Pajarito Plateau encountered coarse-grained deposits that
interfinger with Mio-Pliocene volcanic rocks of the Jemez volcanic field
(WoldeGabriel et al., 2001; Smith, 2004). Deep water-supply wells
drilled beneath the Pajarito Plateau at Los Alamos, New Mexico, encountered gravels beneath the Puyé Formation (Purtymunn, 1995). The
presence of these coarse-grained deposits has been interpreted as an
early axial river system that probably existed during middle or late
Miocene time (Smith, 2004).
Along the eastern margin of the basin are a series of erosional
surfaces preserved beneath gravelly deposits derived the Sangre de
Cristo Mountains (Manley, 1979a, b, 1984; Kelley, 1978). These deposits include the Ancha Formation (Miller et al., 1963; Kelley, 1978;
Spiegel and Baldwin, 1963; Koning et al., 2002), and a suite of progressively inset geomorphic surfaces (with associated deposits) called,
in descending order of age and landscape position, the Oso, Entrañas,
Truchas, Santa Cruz, and Santa Barbara surfaces (Manley, 1979b, 1984).
Ages of these deposits indicate they were deposited coevally with the
volcaniclastic Puyé Formation (Smith, 2004), which was deposited along
the footwall of the Pajarito fault.
Determining ages of incision and the development of the Rio
Grande Valley is complicated by volcanic and structural interactions.
At White Rock Canyon, where the Rio Grande presently flows into the
Albuquerque basin, the ancestral Rio Grande incised at least 250 m
and multiple episodes of deep incision and partial aggradation have
been documented (Reneau and Dethier, 1996). Development of the
present Española Valley of the Rio Grande is constrained by the San
Diego Canyon ignimbrite (1.71 Ma; Turbeville and Self, 1988). This
volcanic ash is near the top of the Puyé Formation, exposed just below
133
a soil-bounded contact with the disconformably overlying 1.62 Ma lower
Bandelier Tuff (Waresback and Turbeville, 1990). The lower Bandelier
Tuff fills a paleocanyon exposed in the walls of White Rock Canyon,
and the upper Bandelier Tuff fills canyons cut in the lower Bandelier
Tuff (Reneau and Dethier, 1996). At White Rock, New Mexico, the lip
of the present canyon contains cobbles deposited by the ancestral Rio
Grande that overlie remnants of the upper Bandelier Tuff. Thus, the
cessation of long-term aggradation in the Española basin probably ended
in early Pleistocene time (Waresback and Turbeville, 1990; Koning et
al., 2002), but the height of the ancestral Rio Grande and its tributaries
have changed (up and down) by more than 250 m during Pleistocene
time. Volcanic ashes derived from the Jemez volcanic field indicate
that streams associated with the southern Sangre de Cristo Mountains
cut into basin-fill of the Ancha Formation between 1.6 and 1.2 Ma
(Koning et al., 2002).
Evidence for earlier (Pliocene) entrenchment by the Rio Grande
also comes from White Rock Canyon, where the Rio Grande excavated
a deep valley into lavas of the Cerros del Rio volcanic field (Reneau
and Dethier, 1996). Pliocene entrenchment in White Rock Canyon has
been interpreted to indicate widespread entrenchment of the entire Rio
Grande fluvial system (Cole et al., 2001; Cole and Stone, 2002). This
early entrenchment, however, could also be the result of the Rio Grande
re-establishing grade during repeated emplacement of Cerros del Rio
lavas on the rising footwall of the La Bajada fault (Connell, 2004; Smith,
2004).
Albuquerque Basin
The Albuquerque basin is the second largest structural basin of
the Rio Grande rift and represents a transitional tectonic feature between the topographically and structurally well-defined northern Rio
Grande of northern New Mexico and southern Colorado, and the broader
Basin and Range to the south. This basin also represents the southernmost part of the contributory section of drainage in the Rio Grande rift.
The northern part of the Albuquerque basin contains the best
direct evidence for Miocene-aged external drainage of the northern Rio
Grande rift (Connell, 2004). Although there are only scant data supporting the occurrence of a single axial river in the Española basin
before Pliocene time, the presence of extrabasinally derived quartzitebearing gravels underlying 7 Ma volcanic rocks exposed along the southeastern margin of the Jemez Mountains indicates that external drainage entered the northern Albuquerque basin since at least late Miocene
time (Fig. 5C; Smith et al., 2001). During the late Miocene, the Popotosa
Formation represented internal drainage at the southern end of the Albuquerque basin (Fig. 7a; Machette, 1978; Connell, 2004).
Extrabasinal fluvial sediments did not reach the southern edge
of the Albuquerque basin until after about 5 Ma when axial-fluvial
deposits of the Sierra Ladrones Formation were laid down (Fig. 7b;
Machette, 1978; Connell, 2004). In the Albuquerque area, extrabasinal
deposits interpreted as ancestral Rio Grande are reported in watersupply wells down to at least 365 m below the land surface (Hawley
and Haase, 1992; Hawley et al., 1995; Connell et al., 1998). Extrabasinal
detritus has also been recognized in some of the deeper water-supply
wells in Albuquerque down to at least 975 m below land surface and
may represent deposits associated with early drainage of the Rio Grande
system.
During Pliocene time, continued uplift along the Sandia Mountains and to the south along the Hubbell Spring fault controlled the
location of the axial river, which once flowed within two kilometers of
the uplifting Sandia Mountains (Connell, 2004). By late Pliocene time,
the ancestral Rio Grande shifted west towards the present valley and
by early Pleistocene time, the river began to incise into older basin fill
between 0.7-1.2 Ma (Connell et al., 2001e, f). Incision of the Rio Grande
Valley has been unsteady and is punctuated by the presence of at least
four terrace deposits associated with periods of valley aggradation. These
terraces are middle to late Pleistocene in age and are discontinuously
134
FIGURE 7. Photographs of Santa Fe Group deposits. Angular unconformity between Plio-Pleistocene basin-margin deposits of the Sierra Ladrones (Palomas?) Formation
overlying deformed playa-margin beds of the Miocene Popotosa Formation in San Lorenzo Canyon, Socorro basin (a). Thickly bedded conglomerate and mudstone rip-up
clasts in fluvial deposits of Sierra Ladrones Formation, Albuquerque basin (b). Trough-cross stratified fluvial sediments of Palomas Formation, Palomas basin (c).
Subhorizontally stratified mudstone and sandstone of Fort Hancock Formation, Hueco Bolson (d).
preserved along the valley (Smith and Kuhle, 1998; Dethier, 1999; Smith
et al., 2001; Connell and Love, 2001).
Central Transport and Distributary Section
The central transport and distributary section of the Rio Grande
is characterized by narrow and relatively shallow basins with an axial
river (Fig. 5D-E). These basins include the Socorro, Jornada, San Marcial,
Engle, Palomas, Coralitos, Hatch, and Rincon (Valley) basins (Fig. 4).
The course of the axial river, considered the ancestral Rio Grande, generally flowed as a single river bounded on both sides by relatively small
transverse-tributary drainages. The Palomas Formation constitutes the
dominant lithostratigraphic unit and is recognized between the Palomas
and Socorro basins (Gordon, 1910; Connell, 2004; Mack, 2004). The
Palomas Formation is divided into an axial-fluvial unit associated with
through-going transport by the ancestral Rio Grande, and piedmont units
associated with transverse-tributary drainages that head in local uplands.
These deposits unconformably overlie deformed beds of the Popotosa
and Rincon Valley formations (Fig. 7A), which were interpreted as having been laid down in separate internally drained basins, (Machette,
1978; Gile et al., 1981)
Socorro Basin
The Socorro basin marks the transition between the relatively
wider basins of the northern Rio Grande rift to the narrower basins of the
central transport and distributary sections. This basin was formed as
uplifts broke up the Popotosa basin during late Miocene time (Cather et
al., 1994). The Popotosa Formation represented deposition in a
fluviolacustrine basin-floor setting, particularly in playa lakes. Volcanic
rocks constrain the end of internally drained basin-floor deposition to
after about 7-8 Ma (Fig. 5D; Connell, 2004).
By Pliocene time, axial-fluvial stream deposits associated with
the ancestral Rio Grande entered the Socorro basin (Modratel??, 1970;
Connell, 2004). The age of the earliest exposed axial-fluvial deposits
of the ancestral Rio Grande are constrained by biostratigraphy and dated
volcanic rocks. The 3.73 Ma Socorro Canyon basalt flow overlies ancestral Rio Grande deposits (Chamberlin, 1999; 40Ar/39Ar date, R.M.
Chamberlin, 2004, personal commun). Biostratigraphic data indicate
that this through-going ancestral Rio Grande was present by at least
2.7-3.7 Ma (Morgan and Lucas, 2003). These data support the presence of extrabasinal sedimentation associated with ancestral Rio Grande
during Pliocene time, but it is not clear when axial-fluvial deposits
first entered this basin. The base of the axial-fluvial succession is not
exposed, but water-supply wells indicate that these deposits are as much
as 365 m thick near Socorro, New Mexico (R.M. Chamberlin, 2004,
personal commun.; cf. cross section in McGrath and Hawley, 1987).
Incision of the basin fill is constrained by tephra from exposures
near the town of San Antonio, about 15 km south of Socorro, New
Mexico. These deposits contain volcanic ash and fluvially recycled
pumice correlated to the lower Bandelier Tuff in deposits of the ancestral Rio Grande that were initially interpreted to be inset against older
basin fill (Cather and McIntosh, 1990), but later investigations demonstrated a fault contact between the pumice-bearing deposits and older
basin fills (Cather, 1996; Dunbar et al., 1996). Ash beds with compositions correlated to upper Bandelier Tuff, one yielding an age of 1.22 Ma,
occur within the upper axial Rio Grande units (Dunbar et al., 1996)
suggesting that deposition of the basin-fill succession ended after 1.2
Ma. Incision of the basin fill is recorded by the presence of at least four
inset fluvial terrace deposits associated with former positions of the
Rio Grande; however, dates for the oldest inset terrace deposits have
not been determined (McGrath and Hawley, 1987).
San Marcial, Engle, and Palomas Basins
The San Marcial, Engle, and Palomas basins are axially linked
basins south of the Socorro basin and constitute the southern part of the
central trunk and distributary section. Paleomagnetic work in the
Palomas Formation resulted in the development of detailed
magnetostratigraphic sections that documented the arrival of
extrabasinal sediment by about 4.9 Ma (Fig. 5E; Mack et al., 1993,
1998). This is supported by microfossil and paleomagnetic data from
fluvial deposits near Truth or Consequences, New Mexico, indicating
deposition between 4.1-4.2 Ma (Repenning and May, 1986). The
Palomas basin contains one of the few places in the Rio Grande rift
where the basin-margin and axial-fluvial deposits are well exposed.
These deposits have been extensively studied by G.H. Mack, W.R.
Seager, and colleagues (e.g., Hawley and Kottlowski, 1969; Hawley et
al., 1981; Gile et al., 1981; Mack and Seager, 1990, 1995; Mack et al.,
1993, 1994a, b, 1996, 1998; Seager and Hawley, 1973; Seager and
Mack, 2003). Stratigraphic and structural studies of the Palomas Formation by Mack and Seager (1990) indicate three stages of basin filling in the region. The preservation of at least 150 m of extrabasinally
derived fluvial sediment near the front of the Caballo Mountains suggests that a major pulse in deformation occurred during Pliocene that
allowed basins to aggrade and move the axial-river towards the basin
master fault (Mack and Seager, 1990). This synorogenic stage was preceded by a long period of internal drainage that was controlled by movement of major basin-margin faults where thick Miocene sediments were
preserved (Mack et al., 1994b). The third post-orogenic stage is interpreted from progradation of tributary deposits basinward from the master
fault (Mack and Seager, 1990).
The age of the upper part of the Palomas Formation is also constrained by early Pleistocene fallout tephras that originated from the
Jemez Mountains (Seager and Mack, 2003; Mack, 2004). Volcanic units
include Plio-Pleistocene fluvially recycled pumice, presumably derived
from the Jemez Mountains, associated with flood deposits in the ancestral Rio Grande facies (Mack et al., 1996). The top of the Palomas
Formation is constrained by the 0.78 Ma Jornada basalt flow that overlies axial-river deposits at the northern end of the Jornada basin
(Bachman and Mehnert, 1978).
Upstream in the San Marcial basin, the ancestral Rio Grande
spilled over and briefly occupied the northern part of the adjacent Jornada
basin. The timing of incision in these basins is poorly constrained because inset fluvial terrace deposits associated with former positions of
the Rio Grande are present in this area have not been dated (Hawley,
1978; Lozinsky, 1985).
Southern Distributary Section
The southern distributary section is characterized by relatively
small tributary drainages within laterally and longitudinally linked
extensional basins (Fig. 4). Distributaries begin their lateral extensions from the area around Hatch, New Mexico, southward. The Camp
Rice and Fort Hancock formations (Strain, 1966) constitute the two
major Plio-Pleistocene deposits of basins in these sections (Fig. 5F-G).
The Camp Rice Formation consists of axial-river, floodplain, and shallow lake deposits and is well preserved throughout most of the Jornada,
Mesilla, and Hueco basins (Hawley et al., 1976; Mack et al., 1997;
Mack, 2004). Deposits of the Camp Rice Formation are as much as 215
m thick in the north-central part of Mesilla basin, but in most places it is
no more than 105 m thick (Hawley and Kennedy, 2004). Broad braided
rivers of the ancestral Rio Grande dominated depositional environments
135
of the Camp Rice Formation (Fig. 7c; Strain, 1966; Stuart and Willingham,
1984; Mack, 2004). The Camp Rice Formation unconformably overlies
the Miocene Rincon Valley Formation and overlaps, and locally
interfingers with, the Fort Hancock Formation (Hawley and Kennedy,
2004). The Rincon Valley Formation is well exposed in the Rio Grande
Valley between the Palomas and Mesilla basins (Seager and Hawley,
1973; Seager et al., 1971, 1975). The Rincon Valley Formation records
deposition within internally drained basins during Miocene time (Seager
et al., 1975; Mack et al., 1998; Mack, 2004).
An extensive lacustrine facies of the Camp Rice Formation is
present in the southern Jornada basin (Lake Jornada of Gile, 2002).
Gypsum is a cementing constituent in deposits of this long-lived lacustrine system of late Pliocene and early Pleistocene(?) age.
Fluviolacustrine environments were primarily fed by distributaries of
the ancestral Rio Grande and the Jornada Draw drainage. Differential
displacement along the Jornada fault zone ultimately produced uplift
of the Tortugas-Doña Ana and Tonuco blocks and topographic closure
of the southern Jornada basin. The ancestral Rio Grande was diverted
to the area of the present Rincon and upper Mesilla valleys during
Pleistocene time (Mack et al., 1997; Mack, 2004).
The Plio-Pleistocene Fort Hancock Formation is exposed in the
southern Hueco Bolson (Hueco structural basin) and is recognized in
water-supply wells in the Mesilla basin (Hawley and Lozinsky, 1992;
Hawley and Kennedy, 2004). The Fort Hancock Formation contains
mudstone and sandstone (Fig. 7d) and records the transition from internal drainage to through-going drainage (Strain, 1966; Hawley et al.,
1969; Seager and Hawley, 1973; Seager et al., 1971, 1982, 1987; Hawley,
1978; Gile et al., 1981; Stuart and Willingham, 1984; Gustavson, 1991;
Collins and Raney, 1994; Mack, 2004). The age of the Camp Rice/Fort
Hancock contact is constrained to about 2.3 Ma by paleomagnetic studies (near the Gauss-Matuyama chron boundary), and by the presence of
the 2.06 Ma Huckleberry Ridge ash (Izett and Wilcox, 1982; Vanderhill,
1986). Thus, integration of the ancestral Rio Grande through western
Texas, as marked by the beginning of Camp Rice deposition probably
occurred in the early part of the late Pliocene.
In the southern Mesilla basin, the age of the upper Camp Rice
Formation is constrained by vertebrate biostratigraphy (Blancan and
Irvingtonian NALMA, Pliocene and early Pleistocene; Fig. 2) and by
tephra (Mack, 2004, and references therein). The 0.76 Ma Bishop ash
is interbedded with uppermost Camp Rice deposits in the lower Mesilla
Valley (Kelley and Matheny, 1983). Just north of the Mesilla basin,
near the Grama siding, the Bishop ash is slightly inset against the Camp
Rice Formation (Seager and Hawley, 1973). The 0.78 Ma Brunhes/
Matuyama polarity chron boundary lies just below the constructional
surface (lower La Mesa, e.g., Gile et al., 1981) that marks the end of
Camp Rice deposition (e.g., Vanderhill, 1986; Seager and Mack, 2003;
Mack, 2004). The stratigraphic position of Bishop ash localities indicates that the end of Camp Rice deposition may be slightly diachronous
(Fig. 5).
The transition from internal to fluvially integrated (external)
drainage in the Mesilla and Palomas basins occurred near the beginning of Pliocene time, perhaps about 5 million years ago (Fig. 5E-F;
Leeder et al., 1996a, b; Mack et al., 1998). During much of Pliocene
time, drainage of the ancestral Rio Grande terminated in the southern
New Mexico border region in a system of interlinked playa lakes that
has been collectively called Lake Cabeza de Vaca (Strain, 1966, 1971).
The maximum elevation of this terminal bolson system is not well constrained, but is estimated to range from about 1130 to 1230 m
(Gustavson, 1991; J.W. Hawley, unpubl. data). In the Hueco basin, the
boundary with the overlying Camp Rice Formation is gradational to
disconformable and local relief on this unconformity may be more than
130 m (Gustavson, 1991). In the Mesilla basin, Camp Rice-Fort Hancock
contact is buried. Studies of water supply wells suggest that the Camp
Rice/Fort Hancock boundary interfingers and rises (in elevation) to the
south (Hawley and Kennedy, 2004).
136
Axial-river deposits of the Camp Rice Formation are also recognized in parts of the Mimbres, Jornada, and Tularosa basins, and smaller
adjacent basins, suggesting that the axial-river periodically spilled over
low-relief topographic divides into adjacent basins (Mack et al., 1997).
Paleosols also appear to be more common in these stratigraphic sections than in sections to the north (cf. Mack et al., 1994a), which suggests that deposition by the ancestral Rio Grande was more sporadic
than in the north.
During late(?) Pliocene time, braided distributary channels of
the Camp Rice axial-river system spread southward and eastward across
Fillmore Pass, south of the Organ Mountains, and ultimately terminated in the extensive playa-lake plains of the Bolson de Los Muertos
of northern Chihuahua, and the Tularosa basin and Hueco Bolson of
New Mexico and western Texas (Hawley, 1975; Hawley et al., 1969;
Strain 1971; Gile et al. 1981; Seager 1981; Seager et al. 1987; Gustavson
1991; Mack et al. 1997). Medium- to coarse-grained fluviodeltaic deposits continued to accumulate on the Mesilla basin floor through early
Pleistocene time (Mack et al., 1998).
Incision of the basin fill is constrained by the presence of the
Lava Creek B ash within the oldest inset fluvial-deposits exposed in
Selden Canyon and at the El Paso Narrows at about 76-91 m above the
present Rio Grande floodplain. This demonstrates that initial cutting
of the Mesilla Valley occurred before 0.64 Ma (e.g., Gile et al. 1981;
Dethier 2001).
The Tularosa basin appears to be a complexly faulted full-graben; however, outcrop, gravity, and well data document two half grabens that tilt away from intrabasinal axial horst (Seager et al., 1987;
Koning, 2002). It is bounded by normal faults of the San Andres and
Organ Mountains to the west and Sacramento Mountains to the east.
The Tularosa basin connects structurally with the Hueco basin to the
south. Little is known about the extent of basin fill, but gravity data
suggest about 2-3 km of Neogene basin fill (Seager and Brown, 1978).
Northern Tularosa basin deposits include alluvial fans and gypsumcemented spring deposits. The southern Tularosa and Hueco basins
periodically received sediment from the ancestral Rio Grande. Projections of the maximum elevation of the Fort Hancock Formation (see
Southern distributary section, see above) northward into the Tularosa
basin suggests that much of this basin was connected to the Hueco
basin and may contain buried Fort Hancock and Camp Rice deposits.
In late Pliocene time, the ancestral Rio Grande spilled through Fillmore
Pass at the southern end of the Organ Mountains and constructed a
large fluvial fan across the basin that created a low topographic divide
separating the Tularosa basin from the Hueco Bolson (Seager, 1981;
Mack et al., 1997). North of the divide are lacustrine and playa-lake
facies of Lake Otero and Lake Lucero, Pleistocene and Holocene lakes
within the Tularosa basin and source of gypsiferous white sands (Lucas
and Hawley, 2002; Langford, 2002; Allen, this volume).
BASINS OF THE UPPER GILA RIVER
The Gila River basin encompasses approximately 145,500 km2
of the Basin and Range physiographic province in the United States
(Fig. 8). The drainage area includes over half of the state of Arizona
and a portion of southwestern New Mexico. The upper Gila River occupies the central part of the Mexican Highland section and parts of
the Mogollon-Datil section of the Transition Zone. The headwaters area
adjoins the Transition Zone and Colorado Plateau provinces on the north.
Surface elevations range from about 3355 m in the Mogollon Mountains of western New Mexico to about 700 m at San Carlos Reservoir,
east of Phoenix, Arizona. The Gila River ends at the junction with the
lower Colorado River at Yuma, Arizona.
The course of the upper Gila River flows through a series of
alluviated troughs outlined principally by north- to northwest-trending
mountain ranges. The river crosses low-lying divides between adjacent
mountain ranges, resulting in the development of bedrock canyons
(knickzones), such as those near Safford, Arizona, and near Redrock,
FIGURE 8. Shaded-relief map depicting the Gila River drainage basin, which
encompasses most of Arizona, and parts of New Mexico and northern Mexico.
Drainages are calculated using DEM and are only shown for reference.
New Mexico. Narrow strips of alluvium exist along the Gila River and its
tributaries. Incised valleys are present at Cliff, Redrock, and Virden,
New Mexico.
Gila Group
Southwestern New Mexico is the extended type area of the Gila
Conglomerate of Gilbert (1875). The Gila Conglomerate was elevated
to group status and subdivided into component formations that range in
age from late Oligocene to middle Pleistocene (Fig. 9; Heindl, 1963;
Krieger et al., 1973; Elston, 1965, 1976; cf. Mack, 2004). The Gila
Group has been mapped in basins of the Transition Zone Province and
Mexican Highland section, west of the Rio Grande rift. Although the
Gila Group was used to describe coarse-grained deposits (e.g., Heindl,
1962; Leopoldt, 1981), it has been expanded to include lithofacies associated with basin-fill environments, ranging from fine-grained lacustrine deposits to fluvial gravel deposits (Ratté et al., 1984; Drewes et
al., 1985). In Arizona the Gila Group is 1-3 km thick and comprises
most of the basin-fill associated with Basin and Range tectonism
(Scarborough and Pierce, 1978; Scarborough, 1989).
The oldest dates near the base of the Gila Group are late Oligocene (~25 Ma). Vertebrate faunas in the upper part are of late Miocene and early to middle Pleistocene in age (Lindsay, 1978; Lindsay
et al., 1984, 1990a; Galusha et al., 1984; Tedford, 1981; Tomida, 1985;
Smith, 1994). Basalts and volcanic ash beds and magnetostratigraphic
correlation support biostratigraphic age determinations (Leopoldt, 1981;
Izett and Wilcox, 1982; Brooks and Ratté, 1985; McIntosh et al., 1991).
The upper part of the Gila Group is largely non-deformed and is mostly
6-3 Ma (e.g., Johnson et al., 1975). Basin-filling probably ended in late
Pliocene or early Pleistocene time (2.5-1 Ma) and was succeeded by
deposition of thin, coarse-grained deposits derived from local uplifts
(Machette et al., 1986). Progressive integration of local drainage systems with the Gila River occurred during Plio-Pleistocene time and
eventually resulted 100-300 m of fluvial dissection (Machette et al.,
1986; Morrison, 1991a).
Much of the emphasis of Gila Group research has been on fossiliferous deposits exposed in the Safford and northern San Simon Valley of Arizona, and in the Duncan and Virden valleys, which straddle
the New Mexico-Arizona state line (Tedford, 1981; Galusha et al., 1984;
Lindsay et al., 1984, 1990a; Morgan and Lucas, 2000). The fine- to
medium-grained deposits in these basins were, for the most part, deposited on broad alluvial plains, and in complexes of ponds and marshes
137
FIGURE 9. Correlation chart showing major stratigraphic units within selected basins of the upper Gila River. Incised deposits (denoted by italics) include terrace (Qt) and
undivided alluvium and eolian deposits (Qa-Qe). Even though the Mimbres basin temporarily received sediment from the ancestral Rio Grande rift, it is placed under
internally drained basins category because for much of its history, it was not linked to the Rio Grande. Data sources listed in appendix.
(cienegas), and ephemeral alkaline playa-lakes. Another component of
this facies assemblage consists of coarse-grained alluvial and debrisflow deposits associated with piedmont slopes flanking adjacent faultbounded range blocks.
Mangas Basin
Mapping at the north edge of the study area in the Mangas structural basin (Leopoldt, 1981) documents final phases of basin aggradation and formation of an integrated upper Gila River system during
Pliocene or early Pleistocene time (Fig. 9A). Deposits of the upper
Gila Group contain mammalian faunas and volcanic ash beds of late
Miocene to late Pliocene age (Leopoldt, 1981; Tedford, 1981; Brooks
and Ratté, 1985; Drewes et al., 1985; Finnell, 1987; Morgan et al.,
1997). Basin floors appear to have been occupied first by ephemeral
saline-alkaline lakes and finally by a system of shallow lakes and
marshes in late Pliocene or early Pleistocene time. Uppermost deposits
of the Gila Group are presently about 215 m above the Gila River floodplain in the central part of the basin. According to Leopoldt (1981) this
lacustrine and cienega system ultimately drained into the Duncan-Virden
sub-basins of the Mexican Highland section through the middle Gila
Box (Redrock, New Mexico) northwest of the Big Burro Mountains.
The uppermost parts of the Mangas and Sapillo Creek drainages probably once drained into the Mimbres basin (Leopoldt, 1981).
Subsequent incision of Mangas basin fill produced stepped valley-border surfaces above the floodplains of the Gila River and its major
tributaries (Leopoldt, 1981). An early(?) to middle Pleistocene deposit
that was part of ancestral Mogollon Creek, and two middle Pleistocene
units are graded to former Gila River base levels from about 150 to 90 m
above the present floodplain. At least two inset fluvial terraces of middle
to late Pleistocene age range from 45 to 20 m above the valley floor.
Gravelly fill of the 20 m terrace can be as much as 20 m thick (Leopoldt,
1981). The youngest valley fill below the Gila River floodplain can be as
much as 30 to 35 m thick (Leopoldt, 1981).
Sapillo Creek is a tributary to the Gila River, northeast of the
Mangas basin. Floodplain deposits of the Sapillo Creek valley are thin
(probably less than 5 m) and the longitudinal valley profile includes
bedrock-controlled knickpoints. Valley border (Sandor, 1983; Sandor
et al., 1986) include a stepped sequence of five surfaces that range
from 8 m to 60 m above the floodplain. These surfaces are cut on lower
Gila Group and are associated with deposits that are less than 5 m
thick. Soils of the four higher surfaces, ranging from 60-12 m above the
valley floor, possess argillic horizons and are considered to be of Pleistocene age (Sandor et al., 1986). A horse skull (Equus conversidens)
collected from one of these terrace deposits is late Pleistocene age (~100
ka, Rancholabrean NALMA, Wolberg, 1980). Well-developed argillic
horizons in soils associated with the two highest surfaces suggest a
middle Pleistocene age (200-400 ka, Sandor et al., 1986).
Higher valley-border surfaces have not been mapped in detail,
but the inferred age of deposits on surfaces about 100 m above Sapillo
Creek is early to middle Pleistocene. Thick gravelly deposits cap the
highest erosion surfaces along Sapillo Creek and the Gila River, respectively, about 8 to 16 km above the confluence with the Gila River. These
surfaces are 250 to 350 m and about 600 m above the floodplain and are
presumably Pliocene in age. They are partly cut on Gila Group and
138
predate canyon entrenchment. Pliocene tephra, present near the eroded
top of the Gila Group, yielded zircon fission-track ages ranging from
2.01-2.13 Ma (unit QTgo of Finnell, 1987).
The San Francisco River, a major tributary of the Gila River that
heads in the northwestern Datil-Mogollon section, now joins the Gila
River in the lower Duncan Valley near Clifton, Arizona. However, the
ancestral upper San Francisco drainage may have previously terminated in the central Mangas basin during Pliocene time (Kottlowski et
al., 1965; Leopoldt, 1981). How and when the upper Gila and San
Francisco rivers integrated with the Safford-San Simon basin and
Duncan-Virden basin remains poorly constrained (cf. Mack, 2004). The
presence of a basalt flow overlying a 90-m high strath at Apache Creek,
about 20 km west of the Continental Divide indicates that incision was
underway by about 1 Ma (Ratté et al., 1984).
Duncan and Virden Basins
The Gila River enters the Mexican Highland section at the southeastern end of the Duncan basin after emerging from canyons through
the Burro Mountains (Fig. 8). The Duncan and Virden basins lie between the Burro and Peloncillo Mountains and straddle the ArizonaNew Mexico border. The Gila River follows a westward course of deep
valleys into the Duncan, Safford, and San Carlos Valleys of Arizona
and intervening canyons cut across major Basin and Range structural
blocks.
Mapping and soil-geomorphic studies of the upper (southeastern) Duncan basin near Virden, New Mexico (Morrison, 1965a, 1991a)
delineated fluviolacustrine facies associated with internally drained
basins (Fig. 9B). Magnetostratigraphy and biostratigraphy constrain
deposition of the upper part of the Gila Group to late Pliocene and
early Pleistocene (3.7 to ~1 Ma; Tomida, 1987; Morgan and Lucas,
2000; Reid and Buffler, 2002).
The oldest gravels in the Duncan and Virden basins (Morrison’s
Unit 6) are locally more than 15 m thick and grade to basin floor levels
that are as much as 140 m above the present floodplain of the Gila River.
These gravels overlie Gila Group deposits (Smith and Mack, 1999; Morgan
and Lucas, 2000; cf. Richter and Lawrence, 1981; Richter et al., 1983),
which are up to 1830 m thick (West, 1996, cited in Reid and Buffler,
2002). The geomorphic position of this unit and strongly developed
petrocalcic soils with Stage IV-V pedogenic carbonate morphology suggest an early Pleistocene or late Pliocene age for this unit. A widespread
gravel and sand (Morrison’s Unit 5) cap basin-floor remnants of Lordsburg
and Pearson Mesas, which are underlain by upper Pliocene Pearson
Mesa Formation of the Gila Group (Smith and Mack, 1999; Morgan and
Lucas, 2000). These mesas are 110-125 m above the present Gila River
floodplain (Morrison, 1965b). The boundary between Plio-Pleistocene
deposits of the Pearson Mesa Member (upper Gila Group) and lower
deformed Gila Group deposits is an angular unconformity (Mack, 2004).
Smith and Mack (1999) interpret the Pearson Mesa Member to have
been deposited in an alluvial-fan and alluvial-flat environment and found
no evidence for an ancestral Gila River in the Virden area (Mack, 2004).
According to Morrison (1985), the Duncan-Virden and Safford
basins were internally drained through early and middle Pleistocene
time and were occupied by large and deep lakes. According to Morrison,
the elevation (above mean sea level) of the highest strand line varies
from 1200 to 1340 m in the Duncan basin and 1130 to 1300 m in the
Safford basin. Morrison considered these differences in elevation to be
the result of subsequent deformation; however the development of a
deep middle Pleistocene lake in the Duncan basin requires more than
100 m of subsequent tectonic deformation in order to account for basin
closure. Repeated closure of the Duncan basin after early Pleistocene
time would presumably involve rather complicated structural deformation to open and close a topographic sill to fill and drain the lakes.
Studies of neotectonic features in the region do not support major Pleistocene deformation (Machette et al., 1986, 1998).
Morrison suggests that establishment of integrated drainage by
the Gila River likely came about because of overflow of the lake in each
basin, instead of by headward erosion and basin capture. This concept
fits well with interpretations of the timing of upstream basin aggradation
and valley incision in the Mangas basin area (Kottlowski et al., 1965;
Morrison, 1965b; Leopoldt, 1981). Morrison’s middle Pleistocene strand
lines could alternatively be interpreted as fluvial-terrace deposits that are
analogous to his (Morrison, 1965a) stream-terraces. Post-early Pleistocene lakes, if present, may not have been nearly as deep or as extensive
as suggested by Morrison (1985). Shallow lakes and cienegas, and lowgradient fluvial fans could just as well have been the major components
of basin-floor depositional environments.
Incision of the Gila River is constrained by the presence of the
0.64 Ma Lava Creek B ash exposed at Nicolas Canyon, a tributary to
the Gila River near Red Rock, New Mexico (Izett and Wilcox, 1982,
NM Site 5; Morrison, 1991a; T. Finnell, 1986, personal communication to J.W. Hawley). This ash is less than 35 m above the Gila River
and apparently sits near the base of fluvial gravels associated with a
channel cut into older basin fill. The height of these terrace gravels
(above local base level) suggest they are inset as much as 60 m below
mesa capping gravels that underlie Pearson Mesa downstream. Thus,
incision of the Duncan-Virden basin probably occurred no later than
early Pleistocene time.
Safford and San Simon Basins
The Safford and San Simon basins of eastern Arizona lie downstream and west of the Duncan and Virden basins of western New
Mexico. The Safford basin contains as much as 4600 m of basin fill
(Kruger, 1991). Before late Pliocene time, the Safford basin was internally drained and contained lacustrine deposits (Fig. 9C; Morrison,
1991b; Houser, 1990; Houser et al., 2002). These deposits contain tephra dated at 2.17 and 2.67 Ma by fission-track dating (Dickson and
Izett, 1981; Galusha et al., 1984). By about 2 Ma (Fig. 9C), throughflowing drainage associated with establishment of the Gila River system is recorded with the transport of extrabasinal detritus through a
gap between the Gila and Peloncillo Mountains called the Gila box
(Houser et al., 2002). Incision of the Gila River in the Safford basin is
recorded by the presence of the 0.64 Ma Lava Creek B ash in the oldest
of a suite of five inset terraces associated with former levels of the Gila
River (Morrison, 1991a; Houser et al., 1985, 2002). This ash is about
24 m above the present Gila River (Houser et al., 2002).
Other constraints on the incision of the Gila River are known
downstream of the Safford-San Simon basins. Downstream in the San
Carlos River Valley, a tributary to the Gila River near Globe, Arizona
(Fig. 8), Anderson (1990) reports that incision of this tributary began
shortly after emplacement of the 3.6 Ma Peridot Mesa flow.
INTERNALLY DRAINED BASINS
Basin fill associated with internally drained structural basins
(bolsons) form the bulk of the Plio-Pleistocene depositional record west
of the Rio Grande and south and west of the upper Gila River. Streams
empty into terminal drainage sinks presently occupied by ephemeral
lake plains or playa lakes. Internally drained basins are also present
along the eastern edge of the Mexican Highland and southeastern Sacramento sections and also in the Transition Zone. Only the western and
northern borders of this region have deep valleys and well integrated
drainage systems (Morrison, 1991a; Hawley et al., 2000; Hawley, this
volume).
The Continental Divide in southwestern New Mexico separates
a number of subsidiary basins with ephemeral-lake plains or playa lakes
(Hawley, 1975, 1993; Kennedy et al., 2000). Basins with large playa
lakes (Animas, Cloverdale, Los Moscos, Playas, Los Muertos, and Wilcox)
were sites of permanent lakes during pluvial intervals of the Plio-Pleistocene (Morrison, 1969; Reeves, 1969; Hawley, et al., 1969, Hawley,
1993; Schreiber, 1978; Williams and Bedinger, 1984).
Alluvial deposits are by far the largest component of the basin fill
139
FIGURE 10. Schematic chronostratigraphic diagram illustrating inception of through-going drainage in basins of the Rio Grande rift. Unconformities are not shown.
Shaded areas represent deposits that contain abundant mudstone and have been interpreted to represent deposition in a fluviolacustrine environment in internally drained
playa-lake basins. Unshaded areas indicate dominantly fluviatile settings, including those with distinctive axial basin floor drainage, such as the ancestral Rio Grande.
Basins are, from south to north, the Hueco basin (HB), Mesilla (M), Palomas (P), and Engle-San Marcial (E-SM), Socorro (S), Albuquerque (AB), central Española
(CEB), and San Luis (SLB) basins.
in this region. Coalescent piedmont-slopes (bajadas) and basin-floor alluvial plains form extensive constructional surfaces. Because there is
little dissection of these basins, deposits have not been extensively studied. Thick lacustrine sequences are also major constituents of the basinfloor facies assemblage at the six playa localities above (Fig. 9D-F).
Closure of lake basins has been produced by three basic mechanisms.
Local structural subsidence of one basin segment relative to adjacent
range and basin blocks is likely the ultimate formative process; however,
closure could be enhanced by a combination of alluvial damming of basinfloor segments by progradation of tributary streams or alluvial fans, and
local deflation of former lake plains with resultant playa-lake floors and
transport of eolian sediment downwind to form local topographically
high areas.
The largest pluvial lake in the region, Lake Palomas (Reeves, 1969)
is in the Bolson de los Muertos-Laguna Guzman area of Chihuahua,
Mexico. It is in a set of deep structural basins that receive drainage from
the distal parts of the Casas Grandes, Santa Maria, Carmen, and Mimbres
fluvial systems, which head into the Sierra Madre Occidental and the
southeastern part of the Transition Zone. The Bolson de los Muertos
was also a temporary base level for the ancestral upper Rio Grande (see
Mimbres basin below). Castiglia (2002) described six latest Pleistocene
through Holocene beach ridges with radiocarbon ages of about 8450 to
220 yrs. BP. A late Pleistocene shoreline shows that small playas were
connected in a larger lake with a surface area of about 7,000 km2. Castiglia
also described a 17-m long core from Laguna El Fresnal that reveals wet
and dry conditions over the past 70 ka.
Important, but areally limited parts of the depositional record
include gravelly veneers on proximal piedmont slopes that cap pediments cut on older basin fill. These unconformity-bounded deposits are
commonly less than 30 m thick, and extend up valleys and canyons of
adjacent ranges as stepped sequences of piedmont-terrace deposits. The
presence of these laterally extensive, unconformity-bounded gravelly
veneers suggests long-term base-level stability and tectonic quiescence
(Menges and Pearthree, 1989).
Neotectonic features include Pleistocene fault scarps in the
Deming-Mimbres, Animas-Lordsburg-Duncan, and San Simon structural
basins (Machette et al., 1986, 1998; Scarborough et al., 1986; and Vincent
and Krider, 1998) and vents and flows in the San Bernadino (Geronimo),
Animas, Potrillo, and Palomas volcanic fields. The Hachita and Playas
Valleys between the Deming and Animas basins appear to have been
tectonically stable since late Miocene or early Pliocene time. Numerous
episodes of volcanism occurred that covered over half of the floor of the
San Bernardino basin by basalt flows. K-Ar dates of rocks of the San
Bernardino volcanic field range from 3.3 Ma on the western flank of the
basin (K-Ar date from Lynch, 1978), to 0.32 Ma (40Ar/39Ar date from L.
Peters, 2004, personal commun.). The presence of the youngest lava
flow near the elevation of the modern floor suggests that little erosion has
occurred since middle Pleistocene time.
Plains of San Agustin
The Plains of San Agustin is a large, mostly non-dissected intermontane basin west of the Rio Grande rift (Fig. 4). The Plains of San
Agustin is one of the few lake basins in the region where shorelines and
related deposits, including wave-cut notches, beaches, bars, and spits,
have been mapped (Weber, 1994). Two cores (200-600 m deep) taken in
the 1950s from the deepest part of the basin encountered older lacustrine
deposits (Fig. 9D; Markgraf et al., 1983, 1984, and references therein).
Early interpretations of the sedimentologic and pollen record indicated
that predominantly fine-grained, unconsolidated deposits extend to depths
of about 290 m and that partly consolidated and conglomeratic fill of
Neogene age is below 326 m (Foreman et al., 1959). Radiocarbon dates
from the upper 9 m from one core indicate a late Pleistocene age near the
top (Markgraf et al., 1983). Paleomagnetic studies of the 164 to 327 m
portion of one core support placement of the Plio-Pleistocene boundary
at depths of 269-305 m below the basin floor (Markgraf et al., 1984;
Hawley, 1993).
Mimbres Basin
The Mimbres basin straddles the border with Mexico, has headwaters in the Black, Diablo, and Pinos Altos ranges in the Transition
Zone, and empties into the Bolson de los Muertos of northern Chihuahua (Fig. 8; Hawley et al., 2000, this volume). The Mimbres basin is
transitional between the Rio Grande rift and internally drained basins to
the west. During late Pliocene time, axial-fluvial deposits of the ancestral
Rio Grande (Cambray fan, figure 6 of Mack, 2004) temporarily emptied
into the southeastern structural margin of this basin (Fig. 9E; Mack et al.,
1997) The basin is bounded on the east by the Goodsight Mountains, the
Sierra de las Uvas, and the basalt flows and ash cones of the West Potrillo
140
FIGURE 11. Generalized paleogeographic diagrams illustrating drainage development in New Mexico and southeasternmost Arizona. During late Miocene time (~10-7
Ma), the northern Rio Grande rift was dominated by through-going rivers that terminated into the southern Albuquerque and northern Socorro basins (a). Internal drainage
dominated basins in southern New Mexico and southeastern Arizona. By early Pliocene time (~5-2.5 Ma), the ancestral Rio Grande drained into southern New Mexico.
Internal drainage dominated persisted in many basins of southern New Mexico and southeastern Arizona (b). By early Pleistocene time (~2.5-0.8 Ma), both the Rio Grande
and Gila River systems drained through the region (c). In much of southwestern New Mexico, internal drainage persisted through most of the Pleistocene.
Mountains. The only major drainage in the Mimbres basin in New Mexico
is the Mimbres River. This river forms terraces upstream and formed a
broad late Pleistocene fluvial fan on both sides of the Florida Mountains
(Seager, 1995; Love and Seager, 1996). The Mimbres fluvial fan terminated into Pleistocene lakes in the Bolson del los Muertos in northern
Chihuahua (Castiglia, 2002).
Eagle Flat Basin
The Eagle Flat basin of Trans-Pecos Texas, between Fort Hancock
and Van Horn, Texas (Fig. 1), is near the boundary between the Sacramento and Mexican Highland sections of the Basin and Range. This small
extensional basin had been internally drained since it formed during Miocene time and had never integrated with the Rio Grande. This basin
exhibits little evidence of Plio-Pleistocene tectonism and is considered an
example of deposition within a tectonically quiescent setting. Langford
et al. (1999) studied a number of sedimentary cores drilled in this basin,
the deepest of which is 219 m below the land surface.
Magnetostratigraphic correlation of selected cores document nearly continuous sedimentation between 12 Ma and middle Pleistocene time (~300
ka). Deposits progressively filled the basin, decreased relief, and eventually onlapped previously active basin-margin structures. By Pliocene
time, the basin was tectonically quiescent. Sedimentation continued until
about middle Pleistocene when the basin became slightly dissected. Sediment accumulation rates generally declined from about 20-13 m/Myr
during Miocene time, to about 9 m/Myr during late Pliocene time
(Langford et al., 1999).
DRAINAGE DEVELOPMENT
Large-scale patterns of drainage evolution in the southeastern Basin
and Range are illustrated by a series of paleogeographic maps based on
interpretations of geologic maps and sedimentologic and stratigraphic
studies, including regional facies patterns, paleocurrents, and sediment
provenance (Figs. 10, 11). Major sources for these maps are discussed in
Connell (2004), Mack (2004), and Smith (2004).
With the possible exception of the Northern San Luis basin, the
transition from internal-drainage to fluvially integrated (external) drainage began earlier in the north and progressively developed southward
(Fig. 10; Bryan, 1938; Hawley et al., 1969, 1976). Structural linkage
between relatively deep basins of the Rio Grande rift in southern Colorado and northern New Mexico required that the northern basins fill
before through-going drainage to the south could be established. Relatively high sediment discharges associated with drainage from high-elevation ranges of the San Juan and Sangre de Cristo Mountains in northern New Mexico and southern Colorado probably enhanced this southward integration by filling the northern basins. The presence of rounded
quartzite gravel in Mio-Pliocene deposits, interpreted as ancestral Rio
Grande, suggest that drainage initially came from the ancestral Rio Chama
and not from the northern San Luis basin (Smith, 2004).
The general lack of extensive fluviolacustrine strata north of the
Albuquerque basin suggests that through-flowing drainage of the ancestral Rio Grande was likely established through north-central New Mexico
before Pliocene time (Fig. 11a). Smith et al. (2001) provided evidence for
pre-Pliocene axial-river drainage in the northern Albuquerque basin. This
early Rio Grande fluvial system probably drained into a terminal basin in
the southern part of the Albuquerque basin until early Pliocene time
(Connell, 2004). As basins of the Rio Grande rift filled, drainage became
integrated through topographic gaps between adjacent basins (Fig. 11b).
By early Pliocene time, through-flowing drainage of the ancestral Rio
Grande was established through New Mexico (Mack et al., 1998). By
late Pliocene time, the ancestral Rio Grande flowed into western Texas
(Fig. 11c; Gustavson, 1991; Mack, 2004). Drainage of the uppermost
part of the Rio Grande was modified by the Plio-Pleistocene Taos volcanic field, which likely impounded drainage in the of the San Luis basin.
Stratigraphic evidence supports the presence of internally drained
basins of the upper Gila River during much of Pliocene time. By late
Pliocene or early Pleistocene time, basins of the ancestral Gila River
became integrated to the Safford area, and by early Pleistocene time, this
river system had incised into the Gila Group. Integration of the San
Francisco River with the Gila River system probably occurred during
late Pliocene time and is recorded by the introduction of gravel associated
with Bonito Creek upstream of the Gila box near Safford, Arizona (Houser
et al., 2002). The presence of the Gila River upstream in the DuncanVirden basins is more complicated and somewhat controversial. This
basin might have hosted lacustrine sediments through most of Pliocene
time. If correct, then the integration of the Gila River system through the
Safford-San Simon basin might have occurred through stream capture
upstream of the Gila Box, just east of Safford, Arizona.
Internally drained basins preserved in southwestern New Mexico
apparently have been aggrading during most of Pliocene and Pleistocene
time. For example, the presence of Pleistocene basalt flows near the
present basin floor suggests that widespread deposition ceased before
Pliocene time in the Animas basin of southwestern New Mexico.
By early to middle Pleistocene time, both the Rio Grande and Gila
River had already begun to incise. Also during this time, drainage from
the upper San Luis basin became integrated into the Rio Grande system.
During the remainder of the Pleistocene, episodes of incision and partial
backfilling left behind suites of terrace deposits and valley border fans.
Basins that remained internally drained continued to receive sediment
throughout most of middle and late Pleistocene time.
DISCUSSION
The Rio Grande and Gila River both follow a similar pattern and
direction of drainage growth, although the headwaters regions have different morphologic and hydrologic characteristics (i.e., glaciated vs.
nonglaciated headwaters of the Rio Grande and Gila River, respectively). We propose that growth of the Rio Grande and Gila River systems arose because of a reduction in regional extension rate in the
southeastern Basin and Range. Mio-Pliocene and late Pliocene expansion of drainage did not appear to coincide with major shifts in climate
(Fig. 2). This similarity in timing of regional drainage integration and
the apparent lack correspondence to major climatic shifts does not rule
out climatically induced changes in the rate of basin filling, but it does
suggest that a reduction in basin subsidence rates may have played a
large role in drainage integration where basins aggraded faster than
they were subsiding. It is possible that rapid subsidence of the basin
floor could generate a greater potential for headward erosion and groundwater sapping that would enhance drainage integration; however, the
timing and downstream direction of drainage integration suggests that
integration occurred through? spillover across slowly subsiding basins.
Miocene deposition is considered to have occurred within tectonically
active extensional basins, resulting in the preservation of thick basinmargin conglomerate and basin-floor mudstone successions representing internally drained basinal settings. By the end of the Miocene, extension rates probably slowed and the basins continued filling. Later
deposition during Pliocene and early Pleistocene times occurred within
tectonically quiescent basins (or those with only minor Pleistocene fault
activity; Machette et al., 1998, Pearthree, 1998), which allowed axial
rivers to connect adjacent basins across low-lying topographic divides.
According to Mack and Seager (1990), Pliocene aggradation in
the Palomas basin of southern New Mexico was controlled by tectonic
subsidence. Mack and Seager (1990) proposed two stages of basin aggradation during Pliocene and Pleistocene times that occurred after an
extended period of internal surface drainage. Their Pleistocene postorogenic stage maybe somewhat misleading because the presence of
Pleistocene-age fault scarps and instrumental seismicity in the Rio
Grande rift indicates that extension is ongoing (Machette et al., 1998).
It is likely that the position of the basin-floor facies may reflect ongoing tectonics (cf., Leeder and Gawthorpe, 1987), but to a lesser extent
than during Miocene and Pliocene times.
It is likely that basins of the northern Rio Grande rift filled first
with sediment because sediment accumulation rates are higher for parts
of the northern contributory section. Plio-Pleistocene accumulation rates
range from 20-190 m/Myr for this northern contributory section of the
141
rift (Rogers et al., 1992; Lozinsky, 1994). The northern contributory
section also contains many large tributaries (i.e., Rio Chama, Rio Jemez,
Rio Puerco, and Rio Salado) that can deliver sediment to the axial
river. Filling these basins might have occurred through progressive onlap
of fluvial sediment across older basin fill (e.g., Cather et al., 1994),
although stream capture across low-lying drainage divides could also
create local unconformities as newly integrated drainage adjusts to base
level differences that are likely to arise from such capture events.
There appears to be a lag in the timing of Rio Grande integration
during Pliocene time that might be related to the alignment of structural basins. North of the Socorro basin, the Rio Grande is confined to
long, axially connected extensional basins. To the south, both axially
and laterally aligned basins would allow for episodic shifts in the position of the axial river into nearby basins. As such basins filled, drainage would eventually become fully integrated where it could flow into
southwestern Texas and northern Mexico.
During Plio-Pleistocene time, deposits of the ancestral Rio
Grande in the Albuquerque basin generally coarsen upwards (Connell
et al., 1998). This increase in sediment caliber suggests an increase in
stream competency or availability of coarse-grained detritus, perhaps
in response to increased effective moisture (e.g., Zhang et al., 2001),
slowing of basin subsidence rate (Blair and Bilodeau, 1988; Mack and
Seager, 1990), or expansion of drainage-basin size (Fraser and DeCelles
1992).
A recent compilation of global data documents a 2- to 10-fold
increase in sediment accumulation rates over the past five million years
(Zhang et al., 2001) that cannot be solely attributed to changes in tectonics (Molnar, 2004). Increased sedimentation expected during Plio-Pleistocene time is, at first glance, not apparent for basins of the Rio Grande
and Gila River. Sparse stratigraphic and magnetostratigraphic data for
basins of the Rio Grande rift suggest that the rate of basin filling was
variable. Initial sedimentation was relatively slow during early and middle
Miocene time, but increased considerably during the late Miocene. Stratigraphic data for the Española basin suggests that stratal accumulation
rates for fluviatile deposits of the Chamita and Tesuque formations may
have been on the order of about 100-200 m/Myr during the late Miocene
(McIntosh and Quade, 1995). Accumulation rates as high as 600 m/Myr
were reported for late Miocene fluviolacustrine deposits in the Albuquerque basin (Lozinsky, 1994); however, there is only scant or inconclusive data for such high accumulation rates. Studies of eolian and fluvial
deposits exposed in the northwestern part of the Albuquerque basin
yield stratal accumulation rate estimates of 69-83 m/Myr during middle
Miocene time (17-12 Ma; Tedford and Barghoorn, 1999), and 150-200
m/Myr during late Miocene time (~10-6(?) Ma; Connell, 2004). PlioPleistocene accumulation rate estimates for the Albuquerque, Palomas,
and Mesilla basins are about 20-30 m/Myr since the beginning of Pliocene
time (~5-0.8 Ma; Mack et al., 1993, 1998; Lozinsky, 1994; Connell and
Love, unpubl. data). Stratal accumulation rates were estimated for both
basin-floor (axial-fluvial and fluviolacustrine) and basin-margin (fluvial
fan, tributary stream, and alluvial fan) settings, and do not account for
the total sediment transport because drainage had already become fully
integrated in New Mexico by Pliocene time. Stratal accumulation rates
were not adjusted for compaction. If compaction adjustments are made
to stratal accumulation rates, then finer grained deposits of the middle to
late Miocene succession would result in even higher rates. The generally
coarser grained Plio-Pleistocene successions were not deeply buried, so
accumulation rate estimates for post-Miocene deposits are probably
reasonable.
Rates of sediment accumulation within internally drained, tectonically quiescent basins, such as the Eagle Flat basin remained low
since Miocene time, with values of about 9-20 m/Myr (Langford et al.,
1999). In the tectonically quiescent San Pedro Valley of southeastern
Arizona, which contains deposits associated with both internal and external surface drainage, paleomagnetic studies indicate accumulation rates
of 20-70 m/Myr (Smith, 1994). These low rates of sedimentation might
142
reflect slower rates of basin subsidence so that most of the sediment is
conveyed through basins, rather than stored. If this is the case, then these
Plio-Pleistocene basin fills should reflect transport rather than
aggradational regimes. The relative paucity of fine-grained sediment in
late Pliocene and Pleistocene axial-fluvial deposits suggests that much of
the suspended stream load was transported out of the region, probably
to the deltas of the Rio Grande and Colorado River.
The presence of weakly to non-deformed Plio-Pleistocene sediments over moderately deformed Miocene rocks also suggest that this
later phase of regional fluvial integration throughout the Rio Grande
rift of northern New Mexico to western Texas occurred during or after
deformation slowed. Such unconformities were preserved as sediments
onlapped onto structural margins during periods of reduced basin subsidence and eventually filled to the levels of low-relief topographic
boundaries across basins. This scenario likely resulted in downstream
integration of drainages as one basin filled and spilled out onto lowerlying basins (Mack et al., 1997).
Long-term incision marks the latest phase of Basin and Range
evolution. This shift towards a dominantly erosional mode of sediment
transport was punctuated by periods of aggradation, resulting in the
development of characteristic stepped valley border profiles, where terrace deposits represent periods of aggradation that were punctuated by
periods of deep and progressive incision (Gile et al., 1981; Pazzaglia and
Hawley, 2004). The timing and character of regional entrenchment is
constrained mostly by reasonably well-dated alluvial chronologies where
the timing and distribution of widespread unconformities can be been
documented. Regional incision of the formerly widespread basin fill is
generally constrained by paleomagnetic determinations of the youngest
basin fill and the age of the oldest, clearly inset deposit, which is commonly well constrained by the 0.64 Ma Lava Creek B ash (Dethier,
2001). This transition from relatively widespread aggradation to incision
occurred near the Brunhes/Matuyama geomagnetic polarity boundary, at
about 0.8 Ma (Mack et al., 1998), and may correspond to intensification
of glaciations (cf. Morrison, 1991b; Zachos et al., 2001). Widespread
basin-fill deposition in basins presently integrated with the Gila River
also ceased shortly after about 0.7-0.8 Ma (cf. Menges and Pearthree,
1989; Smith, 1994; Mack 2004).
The presence of multiple terraces along the Gila River and Rio
Grande indicate that incision of the valley was not steady. Gile et al.
(1981) suggested that the Rio Grande valley was formed by repeated
cycles of incision and partial backfilling that were dominated by climatic changes related to continental glaciations. The presence of progressively inset fluvial deposits along the margins of the modern valley
indicates that episodes of prolonged higher discharge were necessary
to flush sediment and erode the valley. Such incisional episodes must
have occurred prior to aggradation of valley-floor and valley-border
deposits and surfaces.
An increase in frequency and amplitude of climatic shifts during
Pliocene and Pleistocene times might have driven incision of the basin
fill. Increases in bed shear stresses in rivers due to increased discharge
would increase the capacity of streams to transport sediment during
glacial and deglacial times. The crossing of sediment entrainment thresholds, when amplified by rapid changes in vegetation composition, would
create conditions where sediment is easily delivered to drainages (Bull,
1991). If the climate oscillates strongly and rapidly between warmerdrier to cooler-wetter modes, the landscape might not respond quickly
enough to any given change (cf. Fernandes and Dietrich, 1997). Vegetation is quite sensitive to moisture flux variations and could have a profound effect on protection of the landscape from erosion. Bull (1991)
proposed that climate controls sediment production and transport during the transition from moist to dry conditions. Schumm (1965) proposed that incision occurs during glacial times and competence and capacity increases such that streams can remove sediment and incise. During deglacial and interglacial times, the trunk streams can no longer transport sediment and aggrade. The timing of terrace formation for both of
these models would be slightly different. Geochronologic data on terrace
formation in the Rio Grande rift tends to support valley excavation
during glacial times and aggradation during deglacial and interglacial episodes (e.g., Gile et al., 1981; Dethier et al., 1988; Dethier and Reneau,
1995; Reneau and Dethier, 1996; Connell and Love, 2001; Rogers and
Smartt, 1996). Discussions on the timing and possible mechanisms of
river incision and terrace formation are presented by Pazzalgia (this
volume) and Love and Connell (this volume).
Molnar (2001) suggested that increased aridity could amplify
erosion and sediment transport in semi-arid environments. Thus, increases in frequency and oscillation of climate in a semi-arid regime
could amplify erosion and sedimentation rates. Paleoenvironmental data
for Neogene deposits in New Mexico favor a semi-arid climate (Tedford,
1981; Tedford and Barghoorn, 1997), so increased sediment production might not be completely attributed to increasing aridity alone without relying on other threshold conditions to be attained that might destabilize catchments.
Fluvial integration across knickzones in the Rio Grande rift locally played important roles in the evolution of fluvial systems. In the
San Luis basin, Wells et al. (1987) proposed that capture and integration of upper Rio Grande by Red River resulted in a major increase in
drainage area that drove incision of the entire fluvial system. The apparent consistency in timing of stream capture of the upper Rio Grande
and regional incision in both the Rio Grande and Gila River systems,
suggest that capture of the Rio Grande alone might not have driving
incision for this river. Climatically induced changes to headwaters hydrology could aggrade the northern San Luis basin, decrease the height
of the drainage divide, and spill over into lower basins that would accelerate knickzone formation.
SUMMARY AND CONCLUSIONS
Stratigraphic, sedimentologic, and geomorphic data for well-studied Neogene intermontane basins of the southeastern Basin and Range
in New Mexico, southeastern Arizona, western Texas, and northern
Mexico document the evolution of drainage of the Rio Grande and Gila
River, as well as smaller streams that flow into internally drained extensional basins. The Gila River drains high, nonglaciated plateaus
and ranges of the Transition Zone and enters the Mexican Highland
section of the Basin and Range. The Rio Grande heads in Southern
Rocky Mountains that were repeatedly glaciated during the Pleistocene.
The ancestral Rio Grande first formed in northern New Mexico and
probably emptied into terminal playa-lake basins in the southern Albuquerque basin until late Miocene time. Afterwards, the ancestral Rio
Grande extended into southern New Mexico where it emptied into a
series of playa lakes in the Mesilla, Coralitos, Jornada, Mimbres,
Tularosa, and Hueco basins. By late Pliocene time (~2.3 Ma), the ancestral Rio Grande flowed southeastward through Texas and northern
Mexico.
The development of the Gila River is not as well constrained as
for the Rio Grande, but the Gila River appears to have integrated with
basins of eastern Arizona during late Pliocene time. By middle Pleistocene time, the Gila River had already begun to incise and form its
valley.
Basins between the Rio Grande and Gila River, as well as those
east of the Rio Grande drainage remained internally drained throughout the Neogene. Many of these internally drained basins have not been
incised and lack exposures to evaluate their geologic history. A few
have been cored and record progressive aggradation of fluviolacustrine
sediments within tectonically quiescent (or very slowly subsiding) basins.
The general down-stream progression of regional drainage integration does not appear to coincide with major climatic events, and
might be a consequence of the overfilling of tectonically quiescent or
slowly subsiding basins and eventual spillover into adjacent basins
across low-lying interbasinal topographic sills. Drainage integration
143
probably occurred during times of decreased basin subsidence, where the
sedimentation rates exceeded subsidence. The apparent lack of angular
unconformities between deposits associated with external drainage and
underlying (internally drained) units in the southern rift support filling
during times of slower basin subsidence. The reduction in both sediment
accumulation rates and tectonic subsidence also suggests that much of
the sediment load was carried through the southern Rio Grande rift by
the ancestral Rio Grande.
Climatic controls on fluvial deposition are suggested by increased
caliber of Plio-Pleistocene axial-river sediments. Basin incision and
the development of river valleys began between 1.2 Ma and 0.7 Ma
and may have occurred during climatic changes in the early Pleistocene.
Episodic incision and partial backfilling of these rivers and tributaries left
behind suites of terrace deposits and valley border fans that were inset
against the basin fill. The mechanisms of climatically induced incision are
not clearly understood, but probably relate to episodes of increased
stream power linked to the increased amplitude and higher frequency of
Pleistocene climatic changes.
ACKNOWLEDGMENTS
This work was supported by the New Mexico Bureau of Geology
and Mineral Resources (Peter Scholle, Director), New Mexico Water
Resources Research Institute, and U.S. Geological Survey. An early
draft of this manuscript was improved by comments by Gary Smith and
Frank Pazzaglia. We are thankful to many of the Native American communities in the Rio Grande rift for their interest in geological work on
tribal lands. We are also grateful to many researchers of southeastern
Basin and Range geology who have discussed various topics over the
years. These researchers include S.M. Cather, R.M. Chamberlin, C.E.
Chapin, D. Dethier, N. Dunbar, T. Finnell (deceased), L.H. Gile, G.H.
Mack, C.M. Menges, W.C. McIntosh, F.J. Pazzaglia, S.L. Reneau, W.R.
Seager, G.A. Smith, C.E. Stearns, and S.G. Wells. William C. McIntosh, Lisa Peters, Rich Esser, Maureen Wilks, and Nelia Dunbar of the
New Mexico Bureau of Geology and Mineral Resources kindly provided isotopic and geochemical geochronology.
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APPENDIX
SOURCES OF DATA FOR CHRONOSTRATIGRAPHIC COLUMNS FOR BASINS OF THE RIO GRANDE RIFT (FIG. 5)
Column A (San Luis basin)
Appelt (1998)
Bauer et al. (1999)
Brister and Gries (1994)
Colman et al. (1985)
Lipman and Mehnert (1975)
Machette (2003)
Pazzaglia and Wells (1990)
Rogers et al. (1992)
Wells et al. (1987)
Column B (Española basin)
Dethier (1999, 2001)
Gonzalez and Dethier (1991)
Dethier et al. (1988)
Galusha and Blick (1971)
Gonzalez (1995)
Koning (2002)
Koning et al. (2002)
Manley (1979a, b, 1984)
Morgan and Lucas (2003)
Purtymun (1995)
Reneau and Dethier (1996)
Smith (2004)
Spiegel and Baldwin (1963)
Statemap (2004)
Stearns (1979)
Waresback and Turbeville (1990)
WoldeGabriel et al. (2001)
Column C (Albuquerque basin)
Connell (2004)
Connell and Love (2001)
Connell and Wells (1999)
Connell et al. (1999, 2001a-f, 2002)
Column C (Albuquerque basin) cntd.
Dethier (1999, 2001)
Love et al. (2001a)
Lozinsky (1994)
Morgan and Lucas (2003)
Smith and Kuhle (1998)
Smith et al. (2001)
Statemap (2004)
Stearns (1953a, b, 1979)
Column D (Socorro & La Jencia basins)
Cather (1996)
Cather et al. (1994)
Chamberlin (1999)
Connell (2004)
Love et al. (2001b)
Machette (1978)
McGrath and Hawley (1987)
Morgan and Lucas (2003)
Newell (1997)
Statemap (2004)
Columns E & F (Engle-Palomas-Mesilla-Hueco basins)
Gile et al. (1981)
Gustavson (1991)
Lozinsky (1985)
Lozinsky and Hawley (1986)
Mack (2004)
Mack et al. (1993, 1996, 1998)
Morgan and Lucas (2003)
Repenning and May (1986)
Column G (Tularosa basin)
Hawley (1986)
Lucas and Hawley (2002)
SOURCES OF DATA FOR CHRONOSTRATIGRAPHIC COLUMNS FOR BASINS OF THE RIO GRANDE RIFT (FIG. 9)
Columns A & B (Mangas, Alma, and Duncan basins)
Drewes et al. (1985)
Love and Seager (1996)
Morgan et al. (1997)
Smith and Mack (1999)
Morgan and Lucas (2000)
Reid and Buffler (2002)
Hawley et al. (2000)
Mack (2004)
Column C (Safford-San Simon basin & San Pedro Valley)
Houser et al. (1985, 2002)
Lindsay et al. (1990a, b)
Morgan and Lucas (2000)
Morrison (1985, 1991a)
Smith (1994)
Tedford (1981)
Column D (Plains of San Augustin)
Hawley (1993)
Markgraf et al. (1983, 1984)
Pazzaglia and Hawley (2004)
Weber (1994)
Columns E & F (SW New Mexico)
Finnell (1987)
Fleischauer (1977)
Fleischauer and Stone (1982)
Hawley et al. (2000, 2002)
Lynch (1978)
Mack (2004)
Morrison (1985, 1991a)
L. Peters (2004, unpubl. data)