The geology around the city of Guanajuato is specially interesting

Transcription

Aranda-Gómez, J.J.; Godchaux, M.M.; Aguirre-Díaz, G.J.; Bonnichsen,Bill; and Martínez-Reyes, Juventino, 2003, Continental edge tectonics of Isla Tiburón, Sonora, Mexico, in Geologic transects across Cordilleran Mexico, Guidebook for the field trips of the 99th Geological
Society of America Cordilleran Section Annual Meeting, Puerto Vallarta, Jalisco, Mexico, April 5–8, 2003: Mexico, Universidad Nacional
Autónoma de México, Instituto de Geología, Publicación Especial 1, Field trip 6, p. 123–168.
123
FIELD TRIP 6: APRIL 5–8
THREE SUPERIMPOSED VOLCANIC ARCS IN THE SOUTHERN CORDILLERA—FROM THE EARLY
CRETACEOUS TO THE MIOCENE, GUANAJUATO, MEXICO
José Jorge Aranda-Gómez1,@, Martha M. Godchaux2
Gerardo de Jesús Aguirre-Díaz1, Bill Bonnichsen3, and Juventino Martínez-Reyes1
INTRODUCTION
The geology around the city of Guanajuato is specially
interesting because of the quality of the outcrops, the
great diversity of the rocks, and the large number of
clearly exposed structures. The objective of this field trip
is to show the participants some of the most outstanding
features in the transitional zone (Figures 1, 2) between
the Mesa Central and the Transmexican Volcanic Belt
(TMVB), with special emphasis on the record of volcanic
activity in the Guanajuato Mining District and nearby
regions. We hope that during this short visit the participants will obtain an understanding of the tectonic and
magmatic evolution of the area, which is, of course,
linked to the complex geologic history of the adjoining
regions.
The field trip will last for four days and consists of
a total of about 27 stops, with variable amounts of time
spent at each one.
The first day of the field trip is spent in the Mining
District. We provide the participants with an overview of
the geologic evolution of the southeastern part of the
Sierra de Guanajuato with the purpose of showing the
lithologies and the formations that are important in the
District. We include in this day one stop at an outcrop of
the pre-volcanic basement, since these Mesozoic rocks
are an important source of clasts both in the early
Tertiary alluvial fan deposits and in the overlying volcanic units.
1Centro de Geociencias, Universidad Nacional Autónoma de México, Campus
Juriquilla, 76230 Querétaro, Qro., México.
@E-mail address: [email protected]
2Department of Geology and Geography, Mount Holyoke College, South
Hadley MA 01075, U.S.A.
Present Address: 927 East Seventh Street, 83843 Moscow ID, U.S.A.
E-mail address: [email protected]
3Idaho Geological Survey, University of Idaho, 83844 Moscow ID, U.S.A.
E-mail address: [email protected]
The second day of the trip we return to the Mining
District to study in more detail the different facies of the
Calderones Formation, which is the principal subject of
our present research in the region. We can see outcrops of
vent structures as well as of proximal, medial and distal
facies of the bedded tuffs.
In the third day of the trip we travel from
Guanajuato to San Miguel Allende, following the boundary between the Mesa Central and the TMVB. Near San
Miguel we can observe the remnants of two large
andesitic volcanoes of Miocene age (~12–10 Ma) and we
visit the scarp of the San Miguel Allende Fault.
Cretaceous marine sediments, strongly deformed in this
locality both by movement along a late Mesozoic or
Paleogene reverse fault and by superposed normal faulting during the Miocene (≥11 Ma).
The morning of the fourth day of the trip the whole
group is taken to see the outcrops of the Mesozoic basement complex of the Sierra de Guanajuato along the road
which goes from the village of La Valenciana to the
Montaña Cristo Rey (Cerro del Cubilete). In the afternoon, those participants who are able to stay with us visit
outcrops of the Comanja Granite, in the central part of the
Sierra de Guanajuato.
PA RT 1 . O V E RV I E W O F T H E
O G Y B E T W E E N Q U E R É TA R O
REGIONAL GEOLAND
LEÓN
PHYSIOGRAPHIC PROVINCES AND MAJOR ROCK GROUPS
In the region located between Querétaro and León two
physiographic provinces of central Mexico come together (Figure 1). These provinces are the Mesa Central and
the TMVB. Observed in detail, the boundary between the
provinces is complex and transitional, as can be seen in
Figure 2. Immediately north of the boundary, outcrops of
GUIDEBOOK FOR FIELD TRIPS OF THE 99TH ANNUAL MEETING OF THE CORDILLERAN SECTION OF THE GEOLOGICAL SOCIETY OF AMERICA
124
ARANDA-GÓMEZ, GODCHAUX, AGUIRRE-DÍAZ, BONNICHSEN, AND MARTÍNEZ-REYES
basal complex is Mesozoic to earliest Tertiary in age and
is made up of rocks of marine origin, metamorphosed and
intensely deformed by shortening; the complex includes
intrusive bodies of diverse compositions and ages (OrtizHernández et al., 1990). The basal complex is exposed in
a narrow belt which trends NW-SE, near the transition
zone between the two provinces (Figure 3). Rocks of this
basal complex are known only in isolated outcrops in
adjoining regions situated to the north (in the Zacatecas
area) or to the south (near the Valle de Bravo and the
northern part of the state of Guerrero). This basal complex has been assigned to the tectono-stratigraphic
province known as the Guerrero Terrane. The Guerrero
Terrane has been interpreted as an island arc and the remnants of the floor of an ocean basin, both accreted to the
rest of Mexico during the later part of the Early
Cretaceous, around 100 million years ago (Tardy et al.,
1991, 1994).
The Cenozoic cover rests discordantly on the basal
complex and consists of continental sediments and sedimentary rocks, which generally occupy topographically
low zones, and subaerial volcanic rocks, which are principally exposed in ranges and higher plateaus. The rocks
of the Cenozoic cover have experienced only extensional deformation and in some places are gently tilted. These
rocks contain the record of the more recent geologic evolution of the region.
Figure 1. (a) Morphotectonic provinces in central Mexico (after
Sedlock et al., 1993). The location of León (L) and Querétaro (Q) is
shown (see Figure 2). (b) Late Cenozoic normal faults of the southern
Basin and Range Province in northern and central Mexico. The most
obvious structures occur north of the Trans Mexican Volcanic Belt
(TMVB). Some faults south of the TMVB have been interpreted as
part of the same province (Henry and Aranda-Gómez, 1992; Jansma
and Lang, 1997).
mid-Oligocene felsic volcanic rocks predominate (Figure
3); they are genetically related to the Sierra Madre
Occidental (SMO) Volcanic Province of western Mexico
(Figure 1). To the south of the boundary, the rocks most
commonly exposed are late Tertiary and/or Quaternary
andesites considered part of the TMVB (Figures 1, 3). On
the basis of their lithologic characteristics, depositional
environments, styles of deformation and ages, the rocks
which crop out between Querétaro and León can be
divided into two great packages, which we refer to informally as the “basal complex” and the “cover rocks.” The
BASAL COMPLEX OF THE SIERRA DE GUANAJUATO
The basal complex (Chiodi et al., 1988; Dávila and
Martínez, 1987) crops out principally in what has been
referred to as the Sierra de Guanajuato (Martínez-Reyes,
1992) and is made up of:
1. Weakly metamorphosed rocks, developed mostly
from original limestones, shales, and sandstones.
2. Submarine lava flows and pyroclastic rocks, dominantly mafic but occasionally felsic (keratophyres),
metamorphosed to lower greenschist facies.
3. Arc-related pre- and syn-tectonic intrusive bodies
which range in composition from ultramafic
(pyroxenites) to felsic (granites, sensu lato).
Diorites and tonalites are by far the most common
rock types in this group; locally they are intruded by
basaltic to andesitic dike swarms.
4. Post-tectonic plutonic rocks (granites, sensu stricto)
with abundant tourmaline.
UNIVERSIDAD NACIONAL AUTÓNOMA DE MÉXICO, INSTITUTO DE GEOLOGÍA, PUBLICACIÓN ESPECIAL 1 GEOLOGIC TRANSECTS ACROSS CORDILLERAN MEXICO
FIELD TRIP 6: THREE SUPERIMPOSED VOLCANIC ARCS IN THE SOUTHERN CORDILLERA—FROM THE EARLY CRETACEOUS TO THE MIOCENE, GUANAJUATO
125
Figure 2. Digital elevation map of the central portion of the TMVB and the southern part of the Mesa Central. The larger volcanoes of the San
Miguel Allende Volcanic Field are shown: P = Palo Huérfano; J = La Joya; Z = El Zamorano; S = San Pedro. Pliocene volcanoes in El Bajío
plain are Culiacán (C) and La Gavia (G). Sierra del Ocote = O. Cities: León (L); M = San Miguel Allende; D = Dolores Hidalgo; F = San
Felipe.
The intense deformation exhibited by the rocks of
Groups (1), (2), and (3), which make up the Mesozoic
portion of the basal complex, is attributed to two periods
of orogenesis. The first pulse of compressive deformation
happened during the latest part of the Early Cretaceous
and is related to the accretion of the Guerrero Terrane to
the North American craton, and the second one is the
Laramide Orogeny of early Tertiary time.
In some places we find resting unconformably on
the eroded basal complex a thick sequence of continental
red beds (i.e., the Guanajuato Red Conglomerate;
Edwards [1955]). In other places mid-Tertiary intermediate to felsic volcanic rocks, genetically related to the
Sierra Madre Occidental Volcanic Province, lie directly
on the Mesozoic basement. In still other places andesitic
rocks related to the TMVB were deposited atop the basement package.
PULSES OF CENOZOIC MAGMATISM
The Cenozoic magmatism in this region took
place in seven distinct pulses (see Figures 4–6):
Pulse 1. This pre-SMO magmatism took place around 51
Ma with the emplacement of the Comanja Granite, a pluton with a present-day exposure some 50 kilometers in
length by 20 kilometers in width and a northwesterly
trend roughly parallel to El Bajío Fault, along which the
range is uplifted.
Pulse 2. This pre-SMO volcanism was a brief episode of
emission of andesitic lavas at 49 Ma (Aranda-Gómez and
McDowell, 1998), contemporaneous with the accumulation of the Guanajuato Red Conglomerate. This magmatism is seen both as subaerial lavas forming packages of
126
Figure 3. Regional geology of the area between Guanajuato (G) and San Luis Potosí (SLP). Other abbreviations: S, Salinas de Hidalgo; SF,
San Felipe; DH, Dolores Hidalgo; L, León; SM, San Miguel de Allende; VM, Veta Madre; AF, Aldana Fault. Note that south of latitude 22°30’
N most of the area is covered by Cenozoic volcanic rocks. Most stratified Eocene fanglomerates and Oligocene volcanics are tilted to the NE.
Inset shows a rose diagram of orientation of the Cenozoic faults in the Luis Potosí and Guanajuato 1:250,000 quadrangles. Sections A-A’ and
B-B’ are diagrammatic and intended only to show the Cenozoic faulting style. Cenozoic volcanic rocks were grouped in a single unit. After
Aranda-Gómez and McDowell, 1998.
several flows each and as small shallow intrusive bodies
which did not quite reach the surface.
Pulse 3. Despite a paucity of radiometric dates on rocks
of this pulse, we consider it to be volcanism that corresponds to an early phase of SMO activity. Centered in
the Guanajuato Mining District, this intense and prolonged period of explosive and effusive volcanism
occurred in early Oligocene time. It includes all units
from the Bufa Ignimbrite, along with its preliminary
pyroclastic surges, mapped as the underlying Losero
Formation, upward through the Calderones Formation
to the andesitic to basaltic lava flows of the Cedro
Formation.
Pulse 4. This series of eruptions occurred around 30 Ma;
it seems to have involved bimodal volcanism, with rather
extensive flows of andesite spatially and temporally asso-
127
Figure 4. Simplified geologic map of the central portion of the TMVB and the southern part of the Mesa Central. The larger volcanoes of the
San Miguel Allende Volcanic Field are shown (PH = Palo Huérfano; LJ = La Joya; EZ = El Zamorano; SP = San Pedro), as well as the fluviolacustrine Río Laja Basin (labeled as Csc (RL)), El Bajío plain is a broad flat area located between León (L) and Celaya (C). Compare regional fault patterns in the Mesa Central and TMVB. Key: SMA= San Miguel Allende; D = Dolores Hidalgo; SF = San Felipe; I = Irapuato; SLP
= San Luis Potosí. Chronostratigraphic units: Csc = Continental sedimentary deposits; Csc(RL) = Fluvio-lacustrine sediments of the Río Laja
Basin; Qba = Quaternary alkalic basalts; Qtpv = Plio-Quaternary andesites; Tv = Tertiary volcanic rocks; Nb = Neogene andesites; Tof =
Oligocene felsic volcanic rocks; K and Ks = Creatceous marine sediments; Mvs = Sierra de Guanajuato basal complex. Modified from OrtegaGutiérrez et al., 1992. ciated with several forms of rhyolite, all with relatively
high silica contents and occasionally with tin and topaz.
These rhyolites were erupted both as flow-dome complexes and as widespread ignimbrites. They appear on the
map produced by Martínez-Reyes (1992) as the
Chichíndaro, Cuatralba and El Ocote Formations. Almost
the entire outcrop area of these three formations is outside the limits of the Mining District (only the
Chichíndaro Formation is exposed within the central part
of the District). These silicic rhyolites, particularly the
Cuatralba Formation and other ignimbrites and felsic
lava flows, cover an extensive region between the
District and the city of San Luis Potosí (Figures 3 and 4).
This magmatic pulse belongs to the peak phase of SMO
volcanism.
Pulse 5. This pulse of volcanism occurred between 27
and 24 Ma, and it is represented in this region by the
large dacitic domes of El Gigante Field (early Miocene),
by widely distributed but not specially voluminous ignimbrites (27–24 Ma), and by Miocene basalts. We consider this pulse as belonging to the late, waning, phase of
SMO volcanism.
Pulse 6. This pulse includes the volcanism that is truly
transitional between that of the Sierra Madre Occidental
and that of the Transmexican Volcanic Belt (Figure 1). It
is manifested as isolated volcanic domes and ignimbrites
of intermediate composition, emplaced/erupted between
16 and 13 Ma (Cerca et al., 2000) and as widespread
andesite flows.
128
Figure 5. Geology of the southeastern part of the Sierra de Guanajuato (simplified from Martínez-Reyes, 1993).
Pulse 7. The events of this pulse took place between 12
and 8 Ma and include the initial products of the TMVB,
represented by broad benches covered by andesite lava
flows and by the earliest dacitic to andesitic stratovolcanoes, located a bit to the north of the main part of this
volcanic province (Figures 3, 4).
HIATUSES BETWEEN PULSES
It is important to point out the hiatuses in magmatic
activity in this region. There was a long hiatus between
the activity associated with the mid-Cretaceous volcano-sedimentary complex and the Tertiary magmatism
in the Sierra de Guanajuato, whose first manifestations
are represented by the Comanja Granite, dated at ~51
Ma (Zimmermann et al., 1990). The andesitic volcanism around 49 Ma (Aranda-Gómez and McDowell,
1998) seems to have followed the emplacement of the
batholith without any important hiatus. After the eruption of these 49 Ma lavas, there was a long hiatus leading up to the eruption of the Bufa Ignimbrite around 36
Ma (Gross, 1975); it was during this epoch of magmat-
ic quiescence that the Guanajuato Red Conglomerate
was deposited. The next important hiatus occurred
between the end of ignimbritic volcanism of the SMO
type at 24 Ma and the volcanism transitional to the
TMVB type at 16 Ma. During this period there was
ongoing deposition of gravels and sands, which gave
rise to the Xoconostle Formation, a fluvio-lacustrine
deposit which filled a broad shallow basin between San
Miguel Allende and Dolores Hidalgo. Fluvial deposits
are still accumulating in the present-day Río Laja basin
(Figure 4). After 16 Ma there has not been any important hiatus in the volcanic activity. In the early part of
this time period (pulse 6, 16–13 Ma), the volcanism was
sporadic and localized; afterward (pulse 7, 12–8 Ma),
the volcanism began to intensify, reaching peak output
between 10 and 8 Ma. Intercalated with all of the volcanic products from 12 to 8 Ma are widespread fluviolacustrine deposits, whose broad distribution in the central part of the TMVB indicates the presence of extensive lake systems contemporaneous with the early to
middle phases of TMVB volcanism (Aguirre-Díaz and
Carranza-Castañeda, 2000).
129
Figure 6. Geology of the Guananjuato Mining District (simplified from Buchanan, 1979).
DETAILED DESCRIPTION OF PULSES
Pulse 1. Pre-SMO magmatism: the Comanja Granite
The Comanja Granite is an intrusive body of possibly
batholithic dimensions, with a surface exposure of
approximately 50 by 20 kilometers. It was initially dated
at 55±4 and 58±8 Ma (K-Ar, biotite: Mugica and
Albarrán, 1983). More recently, Zimmermann and others
(1990) obtained more precise ages of 53±3 and 51±1Ma
(K-Ar, biotite). Like the majority of circum-Pacific
batholiths, it is probable that this body has a range of ages
in its constituent plutons, though these have not been
mapped separately within it. The body is a granite with
abundant K-feldspar (locally as megacrysts several centimeters in length), quartz, biotite and plagioclase, with
textural variations from a coarse-grained core outward to
a more fine-grained marginal facies. Compositional zoning is not immediately obvious in outcrop; however, mineralogical and geochemical studies, which might reveal
the presence of some variety of cryptic zoning have not
yet been carried out. A common phenomenon among
Paleogene calcalkaline granites in other parts of the
Cordillera is the presence of a peripheral ring of very
small tonalitic bodies of slightly greater age and slightly
greater depth of emplacement than the main granite (e.g.,
around epizonal granite bodies in the Eocene Challis system of Idaho [Earl Bennett, personal communication]). It
is possible that detailed mapping of areas around the margin of the Comanja Granite, and/or of the region to the
southwest of Cerro El Cubilete might identify such bodies. The emplacement of the Comanja Granite post-dates
Laramide deformation, since there is no evidence in outcrop of ductile deformation of the granite. As with the
question of subtle zoning, detailed fabric studies might
reveal some effect of the waning phases of Laramide
compression on the mode of emplacement of the granitic
magma, but such studies have not yet been carried out.
The development of the Comanja Granite marks an
important change in the genesis of intrusive rocks of the
Sierra de Guanajuato, because it was formed by magma
relatively rich in potassium, in sharp contrast to the syn-
130
tectonic plagiogranite of Cerro Pelón. Two more characteristics make the Comanja Granite notable; one is the
abundance of tourmaline and the other is the hints of
tungsten mineralization (El Maguey Mine) at its margins,
which is the only known case in central Mexico. The
tourmaline is present as a late magmatic phase disseminated in the granite, as a coarse-grained phase crystallized in radial clusters in small pegmatitic pockets, and as
the ‘cementing’ material in veins both within and beyond
the granite body. A common feature of these veins is the
presence of breccias formed from granite clasts, largely
in jigsaw-puzzle arrangement, with microcrystalline
tourmaline as the matrix. Along the widest shear zones
every style of gradation is present between simple veins
of massive tourmalinite, without granite fragments, and
breccias with large angular fragments of granite separated by veinlets of tourmalinite. The abundance of tourmaline in the above mentioned three paragenetic habits
–magmatic, pneumatolytic and hydrothermal- suggests
an unusually high boron content in the magma, which
could be obtained by some mechanism of pre-concentration of this element. Production and/or contamination of
the magma through fusion of boron-rich sediments, for
example those sediments accumulated in a fore-arc basin
located close to the continent, seem to be one possible
mechanism. An alternative mechanism might be the
operation of a long-lived hydrothermal cell in the roof
rocks above the magma chamber. Another prominent feature of the Comanja Granite is its external ring of polymetallic skarn ore prospects, which are specially common around the southern margin of the granite body.
These skarn deposits suggest two things: leaching of the
metals from the oceanic crust beneath the fore-arc basin,
and emplacement of the batholith at shallow epizonal
depths, 2–4 kilometers below the Paleocene surface, with
concomitant development of a complex system of
mesothermal to epithermal veins.
Pulse 2. Pre-SMO volcanism: 49 Ma andesitic lavas
within the Red Conglomerate These andesites are intercalated with red beds of the
Guanajuato Conglomerate, principally in the lower member of that unit. The most common type of body is packages of subaerial lava flows that were emplaced on an
almost-horizontal surface, on poorly consolidated sediments. In some outcrops there is sparse evidence of inter-
action between the lava and surface water —poorly developed pillows, thin lenses of phreatomagmatic tuffs or
hyaloclastites, and small clastic dikes of red mud which
occupy cracks at the bases of flows that apparently passed
over wet ground. All the features of these flow packages
are consistent with an environment of deposition that
includes alluvial fans and playa lakes. At other localities,
bodies which could be hypabyssal intrusives (sills or
small laccoliths) or invasive lava flows, are found.
Pulse 3. Early SMO volcanism: early Oligocene explosive to effusive volcanism of the Guanajuato Mining
District After a long period of normal faulting and accumulation
of red beds (mid-Eocene to the beginning of the early
Oligocene), a series of voluminous and varied eruptions
began. Although these volcanic rocks presently have a
more or less restricted area of distribution, they are
nonetheless of great importance (Figure 7). This relevance is not only for the geologic evolution of the
District and the Sierra de Guanajuato (and of the entire
SMO province) but also for the later emplacement of the
major economic mineral deposits for which the District is
famous. We do not know how extensively these rocks
may have been deposited originally because most of the
present-day boundaries of the outcrop area are either
faults or stratigraphic contacts with thick deposits of
younger units; however, it is unlikely that they were
deposited in areas far from the District. This pulse began
with the accumulation of the Losero and Bufa
Formations. The first of these two formations is principally made up of subaerial pyroclastic surge layers and of
tuffs of uncertain eruptive style deposited in (and locally
reworked by) shallow water. The Bufa Formation is a felsic ignimbrite with biotite as its mafic phase. This ignimbrite is in general not highly welded, but owing to
moderate welding and extensive and pervasive silicification it is a hard rock which forms prominent cliffs east of
the city of Guanajuato. It locally contains large lithic
clasts of various types; many derived from the pre-volcanic basement.
After the emplacement of the Losero-Bufa
sequence, there was a hiatus of unknown duration, during
which a surface of considerable relief, at least part of
which was erosional, was developed on the poorly welded and poorly silicified top of the Bufa. The time period
Figure 7. Generalized stratigraphic column of the Guanajuato Mining
District. Modified after Buchanan (1979).
for its development, however, needs not to have been
specially long, because syn-volcanic and post-volcanic
faulting produced much of the initial relief. It seems
probable to us that a caldera was formed as a consequence of the Losero-Bufa eruption. There are subtle
indications of a caldera-forming episode in the topography of the region surrounding the District, and in the
regional drainage pattern, which suggest the development of a broad pre-Bufa uplift whose precise form and
relationship to the regional faulting are yet unknown.
This supposed caldera associated with the Bufa ignimbrite must have been profoundly modified by syn-volcanic and post-volcanic normal faulting, by subsequent
volcanic activity, and by erosion, and it may or may not
have been classically circular or oval in form. We conclude very tentatively that the formation of this caldera,
whatever its shape, produced a closed basin in which the
products of the following series of eruptions, those which
produced the dominantly andesitic Calderones and Cedro
Formations, became trapped. Randall and collaborators
(1994) were the first researchers that postulated the idea
of a caldera in the District.
131
The Calderones Formation is a true encyclopedia of
styles of eruption and emplacement of volcanic products.
It includes low- to medium-grade ignimbrites, deposits of
pyroclastic flows of the block-and-ash type, pyroclastic
surge layers related to phreatomagmatic activity, airfall
ash-rich tuffs, minor Plinian pumice layers, lahars, debris
flows, reworked tuffaceous layers deposited in water,
tuff-breccias, and megabreccias. Ubiquitous and characteristic chlorite alteration imparts a green to greenish blue
color to almost all outcrops of the Calderones, suggesting
that almost the entire formation was deposited in bodies
of shallow water, possibly lakes retained inside the (modified) Bufa Caldera. An alternative interpretation of the
alteration may involve processes of hydrothermal circulation through the Calderones tuffs immediately after deposition, even in the absence of lakes. A third style of alteration, propylitic alteration adjacent to veins and dikes, is
of local importance in many outcrops. It is possible that
detailed petrographic and/or geochemical studies of the
mineral assemblages in many parts of the unit might provide a better assessment of the relative importance of syndepositional alteration (lakes), immediately post-depositional alteration (intracaldera hydrothermal cells) and
later post-depositional alteration (propylitic alteration
adjacent to veins and dikes). Because of the alteration,
and also because of the high quantity of accidental fragments, it is difficult to determine with precision the original compositions of the juvenile volcanic products, but
we consider that in general they were andesites and
dacites. Echegoyén (1970) distinguished three members
in the Calderones; in a very rough way, we think that
these members are equivalent to the proximal facies
(lower member), the medial facies (intermediate member), and the distal facies (upper member) of the
Calderones pyroclastic sequence. The source vents of the
Calderones are located just to the northeast of the depositional basin, in a ring dike which crosses the western ridge
of Cerro Alto de Villalpando, and to the north of the basin,
in the Peregrina Dome Field. Given the internal complexity and variety of pyroclastic products in the Calderones,
we consider it possible that other source vents may exist.
Everywhere within the District, the Calderones
Formation passes upward into the Cedro Andesite, which
is a package of lava flows and associated tuffs of
andesitic to possibly basaltic composition. The
Calderones-Cedro transition consists of an interval of
interstratification of very fine-grained green tuffs with
132
dark brown tuffs and isolated lobes of water-affected lava
flows. We think that the Cedro Andesite was fed by a system of dikes with strikes from ENE to NE, which show a
roughly radial pattern on the map of Echegoyén (1970).
Some of the well-exposed dikes that invaded the distal
parts of the Calderones have marginal facies composed
of peperites and isolated crude pillows which provide
evidence of their interaction with shallow waters and/or
with recently deposited wet tuffs. The uppermost layers
of the Calderones resemble phreatomagmatic deposits;
they carry abundant autoclasts and have a stratigraphy,
which is the reverse of adjacent undisturbed portions of
the formation down to the level of the underlying poorly
consolidated top of the Bufa. This phenomenon of
“inverse stratigraphy” is interpreted as the result of
downward coring of the focus of explosive interaction
between the tip of the dike and its host rocks. The Cedro
Andesite passes upward from its base, within a few dozen
meters, from mixed tuffs and lobes of pillowed lavas into
widespread, apparently subaerial, lava flows. In all outcrops that we have seen, even the larger andesite flows
near the base of the Cedro show evidence of interaction
with water; some flows have well-developed spheroidal
weathering (which locally mimics pillows), and the associated pyroclastic deposits contain matrix palagonite.
The Peregrina Dome Complex also belongs to
this third pulse of volcanism. The Peregrina is found
principally in a large dome field in which one can
observe diverse and very complex relationships with the
Calderones Formation. The range of compositions of the
Peregrina complex varies from dacite to fairly high-silica
rhyolite. In the field it is clear that there are layers in all
members of the Calderones which have abundant clasts
of all the lithologies observed in the Peregrina dome
complex. It also appears certain that the youngest
Peregrina domes cut Calderones layers. According to
Echegoyén (1970), there is at least one Peregrina dome
emplaced in the Bufa Formation. Because in the lower
part of the Bufa we find many clasts of a rhyolite with
very delicate flow-banding, there exists the possibility
that the first-erupted domes associated with this pulse
formed a bit before the emplacement of the Bufa or even
contemporaneously with it. We consider it probable that
throughout this entire pulse of activity domes were being
periodically emplaced, and that the formation we call the
Peregrina is diachronous. Thus, it is impossible to establish a unique age relationship between the Peregrina to
the rest of the volcanic units in the District, with the possible exception of the Chichíndaro rhyolite. Chichíndaro
domes and lava flows seem consistently to cross-cut
and/or overlie the Peregrina in the few places where they
are seen in contact. Blind dikes of Chichíndaro are common in some underground workings of El Cubo Mine,
where they can be seen to cut Peregrina rocks (J. José
Reyes-Martínez, personal communication, 2001).
Although we lack geochemical data on the rocks
formed in this pulse, based on the great compositional
changes observed and on the regional tectonic context,
we presently consider as a working hypothesis the following model for the evolution of the magmas involved.
An original body of andesitic to basaltic magma was generated in a subduction zone along the Pacific Coast of
southern Mexico, which dipped to the east or northeast,
with the downgoing slab passing beneath the ancient
Mesozoic suture zone of the Guerrero terrane with the
continental margin. This magma could have caused partial fusion of various low-melting components of this
complex portion of the North American continental crust,
giving rise to a rhyolitic magma with a fairly high
volatile content. Partly as a result of the ongoing regional and syn-volcanic tectonic extension, this second
magma rose to the surface and erupted explosively, producing the Losero and Bufa formations. Slightly later, the
andesitic magma continued its ascent toward the surface,
establishing a shallow magma chamber. Processes of
MASH (melting, assimilation, storage and hybridization)
may have occurred on a small scale, but it seems likely
that the principal process that modified the magma in the
upper part of the chamber was differentiation (also on a
limited scale), which produced dacitic liquids.
Nonetheless, the most voluminous product of the
Calderones and Cedro eruptions was andesite. In some of
its aspects our working hypothesis has a general similarity with the model proposed for the Taupo Ignimbrite of
New Zealand (Freundt et al., 2000).
Pulse 4. Peak phase of SMO volcanism: Sililcic and
andesitic volcanism from 32 to 30 Ma
The andesitic lava flows of the Cedro Formation (sensu
lato) have a broader distribution along the southern
boundary of the Mesa Central than do the other volcanic
units of the Guanajuato Mining District. Cerca and others
(2000) report 32–30 Ma andesites to the southeast of the
District, which they correlate to Cedro. This broad distribution may suggest a drastic change in the style of volcanism, from predominantly explosive (the Losero, Bufa
and Calderones Formations) to largely effusive. Almost
contemporaneous with the outpouring of the younger
Cedro andesites there was an important episode of activity which produced a large number of high-silica rhyolite
domes and flows, along with a lesser volume of ignimbrites of similar composition (Guillermo Labarthe,
personal communication, 2001). These rocks have been
grouped under the name Chichíndaro Rhyolite (e.g.,
Martínez-Reyes, 1992; Cerca et al., 2000). The distribution of this rhyolite is broad and is similar to that of the
Cedro Andesite (s.l.), and various sources have been documented for it, both within and outside of the Mining
District. Cerca and others (2000) dated the Chichíndaro
Rhyolite at about 30 Ma. The Cedro and Chichíndaro
episode could be interpreted as a stage of bimodal volcanism in the Sierra de Guanajuato, which took place
about 30 Ma. In addition to the Cedro-Chichíndaro
domes and flows, voluminous silicic ignimbrites were
emplaced, principally to the north of the District, in the
Sierra de Santa Rosa (where they are mapped as part of
the Chichíndaro Formation by Martínez-Reyes, 1992),
and to the north of the outcrops of the Comanja Granite.
In this latter part of the Sierra de Guanajuato, these pyroclastic rocks have been mapped as the Cuatralba
Ignimbrite, but this large unit in reality is made up of a
series of ignimbrites and intercalated tuffaceous fluviolacustrine sedimentary rocks whose sources have not yet
been determined. Near the city of San Miguel Allende
there are outcrops of El Obraje Ignimbrite, which has a
radiometric (K/Ar) age of ~28 Ma (Pérez-Venzor, 1996)
and which is of a distinctly higher grade than the other
ignimbrites of this region. For this reason we consider it
as the most characteristic example, among the volcanic
rocks seen in this field trip, of the ignimbrites of the
Sierra Madre Occidental. El Obraje Ignimbrite is a thick
unit and one that appears to be very extensive, but its
source is still unknown.
Pulse 5. Waning phase of SMO volcanism: large dacitic
domes of the El Gigante Field, ignimbrites (~24–22 Ma),
Arperos Gabbro, and early Miocene basalts
During the early Miocene there was a change in the volcanism of the region from widespread bimodal volcanism
133
to the formation of large domes of intermediate composition, such as the hills named El Gigante and La Giganta,
and to the emplacement of extensive ignimbrites which
we interpret as the final phases of the SMO volcanism
(24–22 Ma) in the region. Overlying these are basalts
possibly of early Miocene age. This volcanism was all
located outside the Guanajuato Mining District, both to
the northwest and to the southeast of it. There are only
two published reports that describe the Cenozoic geology
of the area around the District, the map of MartínezReyes (1992) and the map of Cerca and others (2000) for
the southeastern portion of the Mesa Central. MartínezReyes (1990) groups several ignimbrite units as the
Cuatralba Formation. However, the youngest of the ignimbrites turn out to have ages between 24 and 22 Ma, as
was found in the sequence of the Mesa San José de
Allende (Cerca et al., 2000), and thus should be considered as events separate from the Cuatralba series of
approximately 30 Ma exposed north of León. MartínezReyes (1990) also documents the presence of a mafic
intrusive body near the town of Arperos, which he called
the Arperos Gabbro. We interpret this rock body as the
subvolcanic equivalent of the early Miocene basalts. The
Cenozoic volcanic sequence to the north of the town of
Arperos is little studied. In that general region are
exposed several ignimbrite units covered by olivine-rich
basalt flows (not yet dated, but probably early Miocene).
Near the town of Arperos it is possible to observe complex contact relationships between the feeder dikes of the
basalts and the enclosing ignimbrites. There are numerous examples of transition zones with intricate mixtures
of both kinds of rock bordering the diabasic dikes and
shallow sills of the Arperos.
Pulse 6. Volcanism transitional between the SMO and the
TMVB: intermediate lavas and domes
In the period between 16 and 13 Ma isolated domes and
lava flows of intermediate composition were formed. In
the San Miguel Allende Volcanic Field (Pérez-Venzor et
al., 1997), close to the volcanoes Palo Huérfano (Figure
9) and La Joya (Figure 10), mid-Miocene andesitic and
dacitic domes have been documented. These include
Cerro Colorado (~16 Ma, K/Ar, biotite: Pérez-Venzor et
al. [1997]) and El Maguey Dome, which underlies the
~10 Ma andesitic stratovolcano La Joya (ValdezMoreno et al., 1998). Cerca and others (2000) also men-
134
tion undated ignimbrites with intermediate composition
(Las Pilas ignimbrite) and 14 Ma andesitic domes just
north of Salamanca, which they interpret as a transitional volcanic stage between the southern SMO and the
TMVB. The andesitic lava flows that crown the high
points of the Sierra de Guanajuato that were mapped as
the Cubilete Formation by Martínez-Reyes (1992) were
dated at 13.5 Ma by Aguirre-Díaz and others (1997). It
is important to note that these same flows (Cubilete
Formation) are exposed both on the upthrown block and
on the downthrown block of El Bajío Fault, which is the
major structure bounding the southern front of the
Sierra de Guanajuato. The age of the Cubilete
Formation provides us with a control on the timing of
the fault.
Pulse 7. Initial products of the Transmexican Volcanic
Belt: stratovolcanoes, domes and lava-capped mesas
During this latest pulse of volcanism, in the period
between 12 and 8 Ma, the styles of eruption and emplacement once again changed definitively. This was probably
related to a change in the type of magma produced after
the re-organization of tectonic plates in the Pacific Coast.
Dominated by andesites, these magmas produced broad
flat lava benches and locally formed the first major volcanoes of the northern part of the TMVB, that are represented by Palo Huérfano (Pérez-Venzor et al., 1997), La
Joya (Valdez-Moreno et al., 1998), and El Zamorano
(Carrasco-Núñez et al., 1979). The TMVB is still active
today, but the active front is located some 150 kilometers
to the south of the San Miguel Allende Volcanic Field
(SMAVF) and of the boundary between the Mesa Central
and the TMVB (Figures 4, 8).
STRUCTURAL PATTERNS
CENTRAL
IN THE SOUTHERN PART OF THE
MESA
The boundary between the Mesa Central and the TMVB
is evident not only in the stratigraphy but also in the
structures. It has been argued (Aranda-Gómez et al.,
1989; Henry and Aranda-Gómez, 1992, 2000) that in this
region we see the true boundary between the Basin and
Range Tectonic Province and the TMVB (Figure 1b).
North of the transitional zone the morphology and structure are controlled by at least two conjugate systems of
faults trending respectively NW-SE and NE-SW (Figures
Figura 8. Approximate limits of the Trans Mexican Volcanic Belt
and general distribution of Miocene-Pliocene and PlioceneQuaternary volcanic rocks in it. Key: SMAVF = San Miguel Allende
Volcanic Field; A = Amealco caldera; NT = Nevado de Toluca; C =
Colima; G = Guadalajara; Mo = Morelia; M = Mexico City; P =
Pachuca; Q = Querétaro; V = Veracruz. After Pérez-Venzor et al.,
1996. Compare orientation of fault patterns with that in the Mesa
Central (Figuras 3, 4).
2-3), whereas to the south structures striking ENE to EW predominate (e.g., Martínez-Reyes and NietoSamaniego, 1990; Pasquaré et al., 1986, 1987a,b, 1988).
In the zone between Querétaro and San Miguel Allende
(Figure 8), both provinces are cut by a less well-studied
system of faults, whose orientations range from N-S to
NNW, called the Taxco-San Miguel Allende System
(Demant, 1978).
The San Miguel Allende Volcanic Field (Figures
3, 8–10) is part of the Cenozoic cover. It is located
between San Miguel Allende and the village of Colón
(Qro), and immediately north of El Bajío depression.
Four larger volcanoes and several smaller centers of
emission, peripheral to the larger ones, stand out in this
volcanic field. The four large volcanoes are the already
mentioned El Zamorano, La Joya, Palo Huérfano and San
Pedro (Figures 3, 8). These peaks form some of the highest elevations in the region. Radiometric ages (K-Ar and
40Ar-39Ar) published at this time vary from ~12 to ~10
Ma. The volcanoes of the San Miguel Allende Field have
morphological features characteristic of volcanic edifices
in a moderately advanced state of erosion. Their morphology contrasts with that of the younger volcanoes,
such as Culiacán and La Gavia (K-Ar ~2.2 Ma; Ban et
al., 1992), situated immediately to the southwest in El
Bajío depression (Figures 2, 3), whose cones are exceptionally well preserved.
135
Figura 9. Generalized geologic map of Palo Huérfano volcano. Key: PH = Palo Huérfano; EP = El Picacho; CC = Cerro Colorado; CE = Cerro
La Elvira; CP = Cerro El Pilón; SMA = San Miguel Allende; CB = Cañada Begoña; Pa = Presa Ignacio Allende; SM = San Marcos; Cal =
Calderón; RR = Rinconcillo; ER = El Refugio; C = Comonfort; J = Jalpilla; LG = Las Gallinas; PC = Peña Colorada; AB = Agua Blanca; OA
= Ojo de Agua; P = Purgatorio; Ja = Jalpa; E = Elvira; DJ = Doña Juana; Ca = Cañajo; A = Alcocer; Es = Estancia. Simplified after PérezVenzor et al. (1996).
The most outstanding morphological features of
the volcanoes La Joya and Palo Huérfano (Figsures 9, 10)
are the great central depressions in their summit areas,
which form semicircular craters with diameters of ~4
kilometers and average depths of 200 to 300 meters.
Their aspect is similar to that of volcanic calderas of the
Galápagos type. However, it is thought that the unusually large size of these depressions (with respect to the
overall sizes of the volcanoes) is the result of differential
erosion associated with intense hydrothermal alteration
around their central vents (Valdez-Moreno et al., 1998).
Lack of caldera-related deposits near the volcanoes support this interpretation.
All of the volcanoes of the San Miguel Allende
Volcanic Field are composed predominantly of andesitic
and dacitic lavas. By reason of their ages and compositions we consider them to be the oldest large-dimension
volcanoes in the TMVB (Pérez-Venzor et al., 1997). The
form of the edifices and the high proportion of lavas relative to pyroclastic products in La Joya and Palo
Huérfano suggest that they may be structures intermediate between stratovolcanoes and exogenous domes.
Typical stratovolcano deposits, such as pyroclastic flows
and tuffs are relatively scarce in both volcanoes (PérezVenzor, 1996; Valdez-Moreno et al., 1998). Like the
active volcanoes in the southern part of the TMVB, the
volcanoes of the San Miguel Allende Field owe their origin to magmatism associated with subduction of the
Cocos plate under the Pacific margin of southern Mexico.
It seems likely that both the rate and the exact direction
136
Figure 10. Simplified geologic map of La Joya volcano (after Valdez-Moreno et al. 1998).
of subduction of the Cocos Plate differ from those of the
earlier subducted Farallon plate that drove the Oligocene
magmatism (Ferrari et al., 1999). These differences may
in turn be responsible for the different bulk compositions
and volatile contents of the Miocene and younger TMVB
magmas, as compared to the earlier SMO magmas.
The stratovolcano Palo Huérfano is located to the
south of the prominent scarp of the San Miguel Allende
Fault, and it covers this structure without being affected
by it (Figure 9). Based on geologic and geophysic data,
Arzate and others (1999) report that this fault continues
to the south buried under younger deposits for over 80
km and passes through Tarimoro and reaches Parácuaro.
In northern flank of Palo Huérfano lavas from this volcano are displaced by normal faults striking approximately N80E (Figure 9); the orientation of these faults
suggests that they are related to the tectonics of the
TMVB (Figures 1, 8). In the highest part of the Río Laja
Basin, in the area bounded by San Miguel Allende,
Dolores Hidalgo and San Felipe, there are extensive
deposits of gravel and sand (Figures 2-4). Near the
Ignacio Allende Dam (Figure 9) these same gravels and
sands are intercalated with lake sediments and with several volcanic units. The precise age of this stratigraphic
sequence which forms the filling of the Río Laja Basin is
uncertain, and the only sites where precise detailed work
has been carried out are to the north of San Miguel
Allende, in Blancan-Hemphillian fossiliferous localities
studied by Carranza-Castañeda (1987). The sediments in
these localities have been dated (fission tracks in zircon
and/or 40Ar-39Ar in sanidine) between 3.5 and 5 Ma
(Kowallis et al., 1998). On the other hand, in the region
around the village of Xoconostle (Figure 3) NietoSamaniego and others (1996) documented similar gravels
intercalated with a rhyolitic ignimbrite whose radiometric age (K-Ar) is ~25 Ma.
We attribute the origin of this fluviolacustrine
basin to the interaction between normal faulting on the
three fault systems of San Miguel Allende, Alcocer-La
Estancia and El Bajío (see Part II of this manuscript) and
the lava flows emitted by Palo Huérfano, which blocked
the outlet of the hydrologic basin in the region presently
occupied by the mouth of the San Miguel Allende
Reservoir. The fossil vertebrate fauna in the Río Laja
Basin deposits is very important, as it is the oldest locality in which it is possible to recognize fossils of animals
which came out of South America mixed with typically
North American fossils (Oscar Carranza-Castañeda and
Aguirre-Díaz, 2001). The migration of faunas between
the two Americas dates from the early Pliocene and is
associated with the closing of the Isthmus of Panama by
volcanic activity.
At the base of the Palo Huérfano Volcanic
Complex (Figure 9) rocks of the basal Mesozoic complex
are exposed (Chiodi et al., 1988; Ortiz-Hernández et al.,
2002), as is the rhyolitic El Obraje Ignimbrite of midOligocene age (K-Ar, 28.6±0.7 Ma). To the west of the
volcanic center occur outcrops of Miocene basaltic
andesites (i.e., the Allende Andesite, K-Ar, 11.1±0.4 Ma;
Figure 9). Lying above them are the products of the Palo
Huérfano stratovolcano.
PART I I . A
BRIEF REVIEW OF THE GEOLOGIC AND TECTON-
IC EVOLUTION OF THE SOUTHEASTERN PART OF THE
DE
SIERRA
GUANAJUATO
GENERAL FEATURES
The Sierra de Guanajuato is an orographic feature that
extends in a continuous manner over a distance of some
80 kilometers, with an orientation N45W. The southwestern front of the range (El Bajío Fault Zone, Figure 5) is
an important boundary which separates two physiographic provinces in central Mexico (Figure 1a). South of El
Bajío Fault (Figures 3, 5, 8) is the TMVB and north of the
fault is the Mesa Central (Aranda-Gómez et al., 1989),
which is considered as an integral part of the Basin and
Range extensional province (Henry and Aranda-Gómez,
1992, 2000). The present-day morphology of the Sierra
de Guanajuato was caused by this Cenozoic extensional
tectonism. Erosion products of the uplifted Sierra were
carried both to the northeast and to the southwest, accumulating in El Bajío depression (Figure 5), at the foot of
the mountains, as well as in the inner part of the Río Laja
Basin (Figures 3, 8). These deposits of gravels, shales
and claystones, all poorly consolidated, contain vertebrate fossil faunas of Pliocene-Pleistocene age, but there
are also some Miocene deposits. Similar deposits are
found in other regions of the southern portion of the
Mesa Central, in the states of Hidalgo, Jalisco and
137
Guanajuato (Carranza-Castañeda et al., 1996). In the
region of San Diego de la Union (Figure 3) these deposits
are covered by flows of Quaternary alkali basalt which
contain mantle xenoliths (Aranda-Gómez et al., 1989).
The rocks exposed in the Sierra de Guanajuato
can be divided into the two great groups referred to: (1)
the basal complex, which includes both Mesozoic rocks
—volcanic and plutonic rocks of the Guanajuato Arc (in
turn part of the larger Guerrero Terrane) and sedimentary
rocks of the Arperos (fore-arc) Basin— and early Tertiary
intrusive rocks (i.e., the Comanja Granite), and (2) the
Cenozoic sedimentary and volcanic cover (Figure 5). The
sequence in the basal Mesozoic complex includes intrusive rocks of different ages (K-Ar, ranging from 157
down to 108 Ma; Ortiz and Martínez-Reyes, 1993) and a
variety of compositions (ultramafic to felsic) and lowgrade metamorphic rocks (derived from volcanic and
sedimentary protoliths of oceanic origin). The basal complex was intensely deformed by two compressive events.
The first took place at the end of the Early Cretaceous
(Neocomian), when the Guanajuato Arc and its accompanying fore-arc (Arperos) basin were accreted to the
North American continent (Figure 12), and the second
occurred during the Paleocene Laramide Orogeny
(Quintero-Legorreta, 1992). The Cenozoic cover package
in the Sierra de Guanajuato consists of Eocene continental red beds and a thick sequence of volcanic rocks of
Oligocene to Miocene age (Figures 5, 6), predominantly
felsic to intermediate in composition, with the earlier felsic rocks having markedly greater volume than the generally later intermediate rocks.
STRATIGRAPHY, AGE DATES, AND GEOLOGIC EVOLUTION
The basal Mesozoic complex and the early Tertiary
Granite
The pre-Tertiary stratigraphy of the Sierra de Guanajuato
consists of two major associations, a volcano-plutonic
association, which comprises the volcanic and sub-volcanic rocks of the long-lived and possibly multiple arc
and its oceanic crustal substrate, and a volcano-sedimentary association, which comprises the sedimentary rocks
(Figure 11) and intercalated tuffs of the fore-arc basin
(Monod et al., 1990; Ortiz et al., 1992; Lapierre et al.,
1992). These rocks range in age from the oldest plutonic
units (Late Jurassic) to the youngest fore-arc basin com-
138
Figure 11. Reconstructed stratigraphy of the Guanajuato Arc and the Arperos Basin. Φ = Tectonic contact (thrust fault). After Ortiz and
Martínez-Reyes (1993).
ponents (Early Cretaceous). A small outcrop of shallowwater limestone (the Albian-Aptian La Perlita Formation;
Martínez-Reyes [1992]) rests discordantly on the metamorphosed Mesozoic rocks north of the city of León
(Figure 3).
Near the Guanajuato Mining District, the volcano-plutonic association (Figure 11) includes: (1) a
thick succession (>1000 meters) of basaltic pillow lavas
and massive submarine flows (K-Ar, 108.4±56 Ma;
Monod et al., [1990]), with relatively scarce basaltic
tuffs; (2) a group of closely-related diabase dikes
emplaced in gabbros (K-Ar, 112 Ma; Lapierre et
al.[1992]), diorites, quartz-diorites and tonalites; (3) a
massive diorite pluton (K-Ar, ~120-122 Ma; Lapierre et
al. [1992]) with hornblende-rich pegmatitic segregations,
locally cut by basaltic dikes; (4) an intrusive body composed of trondhjemite and leucotonalite (K-Ar, ~143-157
Ma; Lapierre et al. [1992]) and other plutons, also intrud-
ed by swarms of diabase dikes. All these units of the
basal complex are found piled one on top of another on a
series of thrusts (Figurres 5, 11). In many outcrops the
thrusts appear almost horizontal and undeformed; however, there are other outcrops in which the thrusts are
clearly folded. In some places the folding of the thrusts
appears to be drag-folding adjacent to Cenozoic normal
faults, while in other places it appears to be unrelated to
Cenozoic structures. In the field, the lowermost unit is
the one containing the pillow lavas (1), which is in turn
covered by the heterolithologic plutonic unit containing
the dike swarm (2), followed by the massive diorite (3),
and then the trondhjemite-leucotonalite (4). Ortiz and
Martínez-Reyes (1993) have proposed an idealized
reconstruction of the original sequence (Figure 11). This
volcano-plutonic association of the Sierra de Guanajuato
has been interpreted as the upper crust of an intra-oceanic volcanic arc (i.e., the Guanajuato Arc of the so-called
139
Figure 12. Early Cretaceous tectonic evolution of the Guerrero terrane. Key: Guanajuato (Alisitos-Teloloapan) arc; BHCP = Central Mexico
Mesozoic Basin; PVSLP = Valles-San Luis Potosí calcareous platform; TG = Guerrero terrane. Lapierre et al., 1991
Guerrero Terrane). If indeed the apparent age range of
these arc-related rocks (157 to 108 Ma; nearly 50 million
years) is real, this portion of the Guerrero Terrane may be
made up of more than one arc system; however, more
detailed dating and geochemical work would be needed
in order to establish or refute this contention.
Even younger ages have been obtained for some
components of the volcano-plutonic association, in the
range of 66 to 108 Ma. In our opinion the value of these
radiometric dates is a bit uncertain, given that most of the
Mesozoic rocks in this area have experienced regional
greenschist facies metamorphism, and some of them
have also been subjected to heating related to the
emplacement of the Eocene Comanja Granite and/or to
hydrothermal activity associated with the mid-Tertiary
volcanism. Any or all of these factors could have modified the argon content of the rocks, rendering the radiometric dates insignificant as to the true age of the
Mesozoic arc magmatism. On the other hand, it is possible that these anomalous young ages actually are signifi-
cant, and that they provide evidence of a second major
period of arc volcanism, following a change in the polarity of the subduction zone (from west- to east-dipping),
before and during the emplacement of the Comanja
Granite. In summary, the volcano-plutonic association
consists of an intra-oceanic arc or arcs of Late Jurassic to
Early Cretaceous age (Ortiz and Martínez-Reyes, 1993).
The volcano-plutonic association described
above is thrust over a volcano-sedimentary sequence,
which in places is strongly deformed. The rocks of this
association are pelagic in character and consist principally of dark laminated limestone and of thin-bedded black
shale, chert, sandstone and siltstone (Figure 11). Pillowed
basalts, hyaloclastites and basaltic tuffs intercalated with
the sedimentary rocks are found at several localities. KAr ages for these volcanic rocks are in the range from 85
to 93 Ma (Cenomanian to Santonian, or Late Cretaceous;
Ortiz and Martínez-Reyes [1993]). As with the similarly
young ages for certain rocks of the volcano-plutonic
association, these radiometric dates also could be some-
140
what untrustworthy. Poorly preserved radiolarian fossils,
recovered by Dávila-Alcocer and Martínez-Reyes (1987)
from rocks of this association give uncertain ages from
Valanginian to Turonian (roughly 135–90 Ma, or Early to
Middle Cretaceous). Nannofossils collected by CoronaChávez (1998) from layers of limestone indicate ages
from Tithonian to Berriasian (roughly 150 to 140 Ma, or
Late Jurassic). On the basis of these data, the volcanosedimentary association is considered roughly contemporaneous with the volcano-plutonic association (Ortiz and
Martínez-Reyes, 1993). An idealized reconstruction of
the volcano-sedimentary association is shown in Figure
11. This association is interpreted as sediments accumulated in a long-lived oceanic basin (the Arperos Basin)
located between the volcanic island arc(s) and the carbonate platforms, which bordered the Mexican subcontinent (e.g., the Valles-San Luis Potosí Platform). The tectonic juxtaposition of the volcano-plutonic association
and the volcano-sedimentary association is thought to be
related to the closing of the Arperos Basin and to the
accretion of the volcanic arc to the southern edge of
North America during mid-Cretaceous time (Tardy et al.,
1991; Figure 12 of this field guide). Neither the original
width of the Arperos Basin nor the distance between the
edge of the continent and the center of the arc is known
for any time period during the long history of the subduction system; however, the Arperos Basin seems to
have had a significant contribution of clastics derived
from the continent, emplaced as turbidites and/or by processes of off-scraping from the downgoing slab. We speculate that these sediments in turn may have served as
sources of potassium and boron during the period of formation of the Comanja Granite magma.
A bit less than 15 kilometers northeast of the city
of León, shallow-water limestones of Aptian-Albian age
(roughly 120 to 100 Ma, or late Early Cretaceous; shown
as the La Perlita Limestone in Figure 5) crop out resting
discordantly on older, metamorphosed rocks of the volcano-sedimentary association (Chiodi et al., 1988;
Quintero-Legorreta, 1992). These limestones include
oolites and calcareous breccias with abundant
ammonites, brachiopods and gastropods. The presence of
Ceritium bustamantii and Psilothyris occidentalis indicates ages of Neocomian to Aptian (roughly 130 to 112
Ma, or Early Cretaceous). A brachiopod (Peregrinella
sp.) indicates a Hauterivian age (132 to 127 Ma, or Early
Cretaceous).
The Comanja Granite (K-Ar ~51±1.3 Ma; Stein
et al. [1993]) is exposed in the core of the Sierra de
Guanajuato, forming a chain of outcrops more than 50
kilometers long (Figure 3). It is emplaced in rocks of the
Mesozoic basement complex after compressional deformation of the region had largely ended. The Comanja is
a medium- to coarse-grained calcalkaline granite with
large subhedral to euhedral; Carlsbad-twinned K-feldspar
phenocrysts set in a generally medium-grained groundmass of light-colored plagioclase, quartz and biotite. The
pluton is in places cut by dikes of pegmatite and aplite
and by a well-developed network of tourmaline veinlets.
Along its contacts the calcareous sediments of the volcano-sedimentary association were transformed to skarns
by contact metamorphism; in some places the primary
textures were completely wiped out. Brittle shear zones
are very common in the vicinity of the intrusive contacts;
some of these zones are present several tens of meters
inward from the margin of the pluton. The textures of
these shear zones are varied, with granite and/or tactite
clasts of varying sizes and angularities firmly cemented
by abundant tourmalinite.
The Cenozoic volcanic and sedimentary cover
Guanajuato Conglomerate
The basal complex is separated by a major angular
unconformity from the overlying formations (Figure 7).
In the Guanajuato Mining District, immediately above
the unconformity, there is a sequence of continental red
beds (1,500 to 2,000 meters thick; Edwards [1955]). This
sequence consists of boulder and pebble conglomerates,
sandstones and siltstones, with sorting that varies somewhat rhythmically from poor to good and bedding that
varies from massive to thinly layered. Near the base of
the sequence there are intercalated andesitic lavas (K-Ar
~49 Ma; Aranda-Gómez and McDowell [1998]). Based
on the lithology of the deposits and the vertebrate fauna,
Edwards (1955) concluded that the Guanajuato Red
Conglomerate consists principally of sediments accumulated in alluvial fans situated at the base of block-fault
mountains that were rapidly uplifted during the midEocene and Early Oligocene. Near the top of the
Guanajuato Conglomerate there is a series of thin finegrained layers with ripple marks and stream cross-bedding, which suggests that by that time the rate of move-
ment on the faults bounding the depositional basin had
diminished. A statistical analysis of the dips of layers in
the stratified units of the cover indicates a complex history of normal faulting. Extension in this area began
around the middle of the Eocene and continued at least
until the later part of the Oligocene (Aranda-Gómez and
McDowell, 1998).
141
banded rhyolite of unknown origin and affinity. The
thickness of Bufa varies from 350 m in the vicinity of the
Las Torres Mine to less than 10 m on Sirena Hill, less
than 5 kilometers from Las Torres. In the southeastern
part of the District, the Bufa ignimbrite has crude columnar jointing, possibly formed during cooling in a zone
that is more densely welded and/or silicified than is typical of the unit.
Losero Formation
Calderones Formation
Resting unconformably on the Guanajuato Conglomerate
lies the Losero Formation, which is a mixed deposit that
passes upward from a well-sorted sandstone with thin to
medium bedding and dark red color to an interval in
which red layers of sedimentary origin are interstratified
with green pyroclastic surge layers. In the upper part of
the Losero Formation, the green pyroclastic surge layers
are predominant. This transitional sedimentary-to-volcanic package indicates the change from a sedimentary
regime to a dominantly volcanic one (Figure 7). The
Losero Formation has been little studied, but it is important in the interpretation of the volcanic sequence that
was deposited on top of it. Its thickness varies from 0 to
55 meters, but it is widespread in the District, generally
being 10 to 20 meters thick. Structures such as ripple
marks, cross-bedding and graded bedding, cut-and-fill
structures and raindrop impressions in the Losero have
been interpreted uniquely as sedimentary features; however, volcanic structures such as very-low-angle (surge)
cross-bedding and accretionary lapilli (?) are equally
common in these deposits, especially in the upper part of
the unit. The depositional environment is interpreted as a
shallow lake (Edwards, 1955), although our observations
suggest that the upper, surge-dominated, layers were
deposited subaerially.
Bufa Ignimbrite
An erosional surface separates the Losero Formation
from the overlying unit, the Bufa Formation (K-Ar,
37>0±3.0 Ma; Gross [1975]), which is an ignimbrite with
less than 25 percent by volume of phenocrysts of quartz,
sanidine and plagioclase, and small euhedral biotite commonly replaced by opaque minerals. Dispersed in the
deposit are lithic clasts of andesite and rhyolite. Near its
base, the ignimbrite contains clasts derived from the
lower units and abundant fragments of a delicately flow-
The Calderones Formation is a complex unit that
includes an indeterminate number of andesitic to dacitic
ignimbrites and layers of volcaniclastic material that
accumulated in a shallow lake. This formation rests on an
eroded and faulted surface developed on the Bufa
Ignimbrite (Figure 7). Calderones commonly fills paleochannels, especially in the proximal and medial areas.
Some of these channels appear to have been formed by
stream erosion prior to the deposition of the Calderones,
while others may have been formed, or at least deepened,
by the passage of the surges and density currents which
gave rise to the basal layers of the Calderones. There are
places where angular lithic fragments of (metamorphosed?) chloritized andesite (derived from the basal
complex?) make up ~75 per cent of the deposit. In other
places fragments derived from sedimentary rocks of the
Mesozoic basement are more abundant than the juvenile
volcanic materials. In still other places, clasts derived
from the growth and/or destruction of one or more domes
in the Peregrina Dome Field (Figure 6) constitute an
important component of the deposit. The Calderones
Formation is medium- to coarse-bedded, and the grain
size of the accidental clasts ranges from fine sand to pebbles and cobbles, although the most common are pebbles
and small boulders. There are some layers with wellrounded clasts, but the majority of the layers has angular
clasts. In the vicinity of El Cubo Mine, Calderones contains pyroclastic flow deposits. These tuffs range from
thin (3 m) to moderately thick (20-30 m), and the base of
each flow is marked by a horizon rich in boulders of purple latite or dacite (probably derived from the Peregrina
Domes) in a vitroclastic and chloritized matrix. Above
this basal horizon there are welded tuffs which display
collapsed pumices that are entirely replaced by chlorite.
In some places there are delicately laminated layers of
fine-grained material that we interpret as pyroclastic
142
surge deposits associated with emplacement of the overlying ignimbrites or as deposits produced in local
phreatomagmatic events. The total thickness of
Calderones Formation (Figure 7) has been estimated
between 200 and 250 m (Buchanan, 1979). However,
there is not sufficient information about the thickness of
the unit in the block bounded by El Cubo and La Leona
Faults for us to be certain that we know the true maximum thickness.
tures as feeder dikes for the Cedro lava flows. There is
evidence that locally the dikes had phreatomagmatic
interaction with the uppermost (generally but not exclusively distal) layers of the Calderones Formation. This
activity produced small lahars that appear to be intercalated with lenses of tuffaceous materials characterized by
their massive nature and by the abundance within them of
andesite clasts derived from the dikes.
Chichíndaro Rhyolite
Cedro Andesite
Resting on the Calderones Formation and interstratified
with its upper layers is the Cedro Formation, which is
made up of grey to black andesite lava flows, in places
with interbeds of pyroclastic materials. The total thickness varies from 100 to 640 m. As in the case of the
Calderones Formation, the age of the outcrops of the
Cedro within the District is not known. Cerca and others
(2000) dated (K-Ar, 30.5±0.5 Ma) an andesite which they
considered as part of the Cedro Formation in the volcanic
sequence of the Mesa de San José de Allende, located
between the cities of Guanajuato and San Miguel Allende
(Figure 3).
Dike Complex
In its lower portion, the volcanic sequence of the District
is characterized by the presence of dikes whose compositions are similar to those of the nearby Cedro andesite
flows. These structures are overwhelmingly most abundant, most elongate and widest where they are exposed
cutting outcrops of the Calderones Formation. Many of
these dikes cut across the La Leona Fault (Figure 6),
which separates surface exposures of Calderones east of
the fault trace from surface exposures of the Bufa west of
it. Invariably these dikes terminate a few meters or tens
of meters into the Bufa block. Elsewhere, Cedro dikes do
persist for long distances in outcrops of the Bufa, especially in the region between the village of Calderones and
the city of Guanajuato, but they are relatively narrow. A
few dikes were mapped by Echegoyén (1970) as cutting
outcrops of Cedro flows, but most dikes are overlain by
these flows. The reason for the paucity of dikes exposed
within the block of Bufa Ignimbrite and Guanajuato
Conglomerate bounded by the La Leona Fault and the
Veta Madre is not entirely clear. We consider these struc-
The youngest volcanic unit in the Guanajuato Mining
District is a rhyolite porphyry that forms large domes,
tholoids and lava flows, along with associated ignimbrites and volcanic breccias. Its type locality is
Chichíndaro Hill (Figure 6), where a large altered dome
of uncertain age and affinities lies between two branches
of the Veta Madre. Somewhat similar volcanic structures
are exposed at the summit of Cerro Alto de Villalpando
in the northeastern part of the Mining District, in other
places to the northeast of the District, and in the northern
part of this end of the Sierra de Guanajuato. In places,
such as the Sierra del Ocote (Figure 2), the rhyolite
domes contain disseminated tin and vapor-phase cavityfilling topaz distributed along the flow foliation. Gross
(1975) reported K-Ar ages of 32±1 Ma for the
Chichíndaro Rhyolite, at its type locality. NietoSamaniego and others (1996) obtained two K-Ar (sanidine) ages on rhyolitic domes that they considered part of
the Chichíndaro Formation, one in the La Sauceda
Graben (Figure 3), south of the District (30.8±0.8 Ma),
and the other north of the town of Santa Rosa (Figure 3),
north of the District (30.1±0.8 Ma).
Cubilete Andesite
At the summit of Cerro del Cubilete, resting directly on
Mesozoic metasedimentary rocks, there is a sequence of
gravel deposits topped by an andesitic lava flow (Figure
5). This gravel, shown on the map of Martínez-Reyes
(1992) as El Capulín Gravel, accumulated in what originally were low zones and/or channels that were the
courses of major streams. Later extensional tectonic
activity is responsible for their presence at an elevation
about 600 m higher than the valley known today as El
Bajío Plain. This gravel is composed principally of small
rounded boulders derived from the mid-Tertiary volcanic
sequence and, in much lesser proportion, of clasts with
provenance in the basal Mesozoic complex. The thick
(presumably intracanyon) andesite(s) produced marked
thermal-alteration effects in the underlying sediments,
and the flow bases are autobrecciated. Sets of vertical
fractures create rough columns in the middle portion of
the flow(s). In the uppermost part of the andesite outcrop,
there is an interval of subhorizontal fracturing, and
between the upper and lower columnar-jointed portions
there is a second surface of platy jointing. This type of
fracture pattern could be developed in a single thick flow
with colonnade-and-entablature structure, or it could
equally well be developed in two thinner lava flows. This
unit is also found in El Bajío Plain, south of El Cubilete,
at an elevation 600 m below the Cubilete summit, where
it commonly forms the tops of benches held up by either
the Cuatralba Ignimbrite or El Capulín gravel. Thus the
andesite, like the gravel (but note that there are broad
areas of much younger gravels, not belonging to El
Capulín, but included with it for mapping purposes, in El
Bajío Plain south of El Cubilete), crops out both on the
upthrown and on the downthrown blocks of El Bajío
Fault Zone. The precise age of the true El Capulín gravels is not known, but the andesite has an age of about 13.5
Ma (Aguirre-Díaz et al., 1997), while the youngest identifiable clasts in El Capulín gravels are those derived
from ~30 Ma ignimbrites. Assuming that the displacement on El Bajío Fault took place between the midMiocene and the present, its long-term rate of vertical
displacement would be on the order of 0.04 millimeters
per year. Of course, the actual period during which fault
movement took place may have been much briefer than
13.5 Ma, but we have no stricter field constraints on it at
present.
Structure
The main structural trend in the southern part of the
Sierra de Guanajuato has a NW-SE orientation, which is
defined by: 1) the schistosity in the low-grade metamorphic rocks, 2) the average attitude of fold axial planes in
rocks of the basal complex, 3) the outcrop pattern of the
Mesozoic basal complex and the Comanja Granite, and
4) the trends of some of the larger Tertiary faults (e.g., El
Bajío Fault and the Veta Madre). Detailed analysis of the
trend and plunge of microfold axes in the metasediments
also shows a less evident NE-SW trend. Tertiary normal
143
faults, such as the Villa de Reyes Graben (TristánGonzález, 1986) and the Aldana fault are also oriented
NE-SW (Figures 3, 5).
The contacts between the principal stratigraphic
sequences (i.e., the volcano-plutonic association
(Guanajuato Arc) and the volcano-sedimentary association (Arperos Basin), and even the contacts between certain lithologic units within each of these two associations
of the metamorphic basement, are persistent subhorizontal mylonite zones of little thickness. Monod and others
(1990) considered these contacts as low-angle thrusts of
mid-Cretaceous age. Overprinted on this first event of
compressive deformation, which caused the formation of
an early foliation, with a distinctive and penetrative
north-south oriented crinkle lineation, is a second compressive deformation of Laramide age (Paleocene-early
Eocene), which produced somewhat larger folds with
axial planes that have northwesterly strikes and northeasterly dips.
The cover rocks display structures associated
with extensional tectonism. There are at least two conjugate systems of faults in the Mesa Central (Figures 2, 3),
which in the Sierra de Guanajuato affect both the basal
complex and the cover. The most important trend in the
Sierra de Guanajuato is the NW-SE trend discussed
above, but there are also important structures oriented
NE-SW and ENE-WSW. Examples of the later are the
Villa de Reyes and La Sauceda grabens, respectively, and
a linear feature between the villages of Los Mexicanos
and Santa Rosa which has been considered a graben by
some authors (e.g., Martínez-Reyes, 1992) and a paleovalley by others (e.g., Guillermo Labarthe, personal communication, 2001). This feature, whatever the origin of its
boundaries is, apparently was filled by early Tertiary conglomerates and mid-Tertiary volcanic rocks (Figure 5).
There is indirect evidence of extensional tectonism in the Sierra de Guanajuato during the Eocene, contemporaneous with deposition of the continental red beds
of the Guanajuato Conglomerate. Aranda-Gómez and
McDowell (1998), based in a statistical analysis, have
suggested that variations in the degree and direction of
dip of both the red beds and the volcanic sequence are
consistent with tilting associated with normal faulting
active during the accumulation of these layers. The
majority of the normal faults has its downthrown blocks
on their southwestern sides; an important exception is the
La Leona Fault (Figure 6), which dips to the northeast
144
and has its northeastern side downthrown. The significance of this exception to the general rule is not yet clear.
There is a possibility that it is related in some way to the
eruption of the Bufa Ignimbrite and the formation of at
least a part of the system of closed basins which a bit later
served to entrap the volcanic materials which make up
the Calderones Formation. The principal northwesttrending faults developed before the mineralization of the
vein systems of the District, because virtually all the ore
deposits were emplaced along these structures. In the
regional setting (the southern portion of the Mesa
Central, Figure 3), it is not possible to establish consistent crosscutting relationships between the NW-trending
faults and the NE-trending ones (Aranda-Gómez et al.,
1989). It is therefore concluded that the pattern of faulting in this broad region cannot be attributed to a single
period of extension in which the orientation of the principal stresses was constant (Aranda-Gómez, 1989). It is
possible that the present-day fault pattern evolved
through successive periods of extension, each taking
place under the influence of a different set of forces
(Aranda-Gómez and McDowell, 1998).
Mineralization
By a considerable amount, the most important mineralization in the Guanajuato District consists of epithermal
veins of silver and gold whose age of formation has been
dated as 27.4±0.4 Ma (Buchanan, 1975). The District has
a history of mining activity of more than 450 years. The
first of the present underground mines was developed by
the Spaniards in 1548, although there is some evidence
that the indigenous peoples of the area had been extracting gold and silver from near-surface deposits for many
years before that date. It has produced ~130 tons of gold
and ~30,000 tons of silver, making it one of the most
important silver districts in the history of precious-metal
mining worldwide. The production is derived from three
principal vein systems (the La Luz, Veta Madre and La
Sierra Systems, Figure 6) of quartz, adularia and calcite,
emplaced both in rocks of the basal complex and in those
of the Cenozoic cover (Figure 6). Concentrations of precious metals are present in isolated packets (known as
bonanzas, or “spikes”) distributed vertically and laterally
between non-mineralized segments of the veins. There
are three principal levels of production at 2,100-2,350,
2,200-1,700, and <1,700 m a.s.l. The mineral associa-
tions in the upper- and middle –level bodies are: acanthite
+ adularia + pyrite + electrum + calcite + quartz. In the
lower-level bodies they are: chalcopyrite + galena +
sphalerite + adularia + quartz + acanthite. This suggests
that the mineralization was produced by fluids of two different compositions (Buchanan, 1979). The veins occupy
what originally were normal faults. The Veta Madre can
be followed on the surface for about 20 km, it dips from
35 to 55 degrees to the SW and it has measured displacements of around 1,200 meters near the Las Torres Mine
and 1,700 meters near La Valenciana Mine.
In addition to the epithermal veins, near
Guanajuato small deposits of stratabound massive sulfides (e.g., Los Mexicanos, Figure 5) have been reported
in the Mesozoic volcano-sedimentary association.
Similarly, there is gold mineralization in the Comanja
Granite, and in its contact aureole small tungsten deposits
have been found. In the Tertiary volcanic rocks, principally in the topaz rhyolites, there are small tin prospects.
Near Cerro del Cubilete, there are tabular bodies where
kaolinite is being quarried. These bodies possibly are
hydrothermally altered Tertiary dikes emplaced in the
basal complex. Finally, the finely laminated Losero beds
and some of the Calderones ash flow tuffs have been
extensively quarried and used for construction (mainly as
a facing stone on large buildings and for construction of
columns and smaller buildings).
ACKNOWLEDGMENTS
Throughout the years our research in the transitional
region between the Mesa Central and the TMVB has
been financed by different agencies. J. Aranda gratefully
acknowledges grants form CONACYT (37429-T) and
DGAPA PAPIIT (INI114198); M. Godchaux gratefully
acknowledges support provided by Mount Holyoke
College for her work in the Sierra de Guanajuato and
adjoining regions; and G. Aguirre thanks CONACYT and
DGAPA PAPIIT for financial support through the grants
33084-T and IN-120999, respectively. BIBLIOGRAPHICAL
REFERENCES
Aguirre-Díaz, G.J.; Nelson, S. A.; Ferrari, Luca; and López-Martínez,
M., 1997, Ignimbrites of the central Mexican Volcanic Belt,
Amealco and Huichapan calderas (Querétaro-Hidalgo), in
Aguirre-Díaz, G.J.; Aranda-Gómez, J.J.; Carrasco-Núñez, G.;
and Ferrari, Luca, eds., Magmatism and tectonics of central and
northwestern Mexico - A selection of the 1997 IAVCEI General
Assembly excursions: México, D.F., Universidad Nacional
Autónoma de México, Instituto de Geología, Excursión 1, p.
1–39.
Aguirre-Díaz, G.J., and Carranza-Castañeda, Oscar, 2000, Las
grandes cuencas del Oligo-Mioceno del centro de México:
GEOS, v. 20, p. 301.
Aguirre-Díaz, G.J.; Zúñiga, R.; Pacheco, F.J.; Guzmán, M.; and Nieto,
Jorge, 1999, El graben de Querétaro, México. Observaciones
de fallamiento activo: GEOS, v. 20, no. 1, p. 2-7.
Aranda-Gómez, J.J.; Aranda-Gómez, J.M.; and Nieto-Samaniego,
A.F., 1989, Consideraciones acerca de la evolución tectónica
durante el Cenozoico de la Sierra de Guanajuato y la parte
meridional de la Meseta Central: Universidad Nacional
Autónoma de México, Instituto de Geología, Revista, v. 8, no.
1, p. 33–46.
Aranda-Gómez, J.J., and McDowell, F.W., 1998, Paleogene extension
in the southern Basin and Range Province of Mexico; syndepositional tilting of Eocene Red Beds and Oligocene volcanic
rocks in the Guanajuato Mining District: International Geology
Review, v. 40, p. 116–134.
Arzate, J.A.; Aguirre-Díaz, G.J.; and Arroyo, M., 1999, Mediciones
geofísicas aplicadas al estudio de la falla Tarimoro-San Miguel
de Allende (SMA); una posible discontinuidad mayor en el
basamento: GEOS, v. 19, p. 237.
Ban, M.; Hasenaka, T.; Delgado-Granados, Hugo; and Takaoka, N.,
1992, K-Ar ages of lavas from shield volcanoes in the
Michoacán-Guanajuato volcanic field: Geofísica Internacional
(Mexico), v. 3, p. 467-474.
Branney, M.J., and Kokelaar, P., 1992, A reappraisal of ignimbrite
emplacement; changes from particulate to non-particulate flow
during progressive aggradation of high-grade ignimbrite:
Bulletin of Volcanology, v. 54, p. 504-520.
Buchanan, L.J., 1979, The Las Torres mine, Guanajuato, Mexico; ore
controls of a fossil geothermal system: Golden, Colorado
School of Mines, Ph.D dissertation, 111 p. (unpublished).
Carranza-Castañeda, O., and Aguirre-Díaz, G.J., 2001, Índices bioestratigráficos de las cuencas sedimentarias del Terciario tardío
del centro de México: Reunión Anual de la Unión Geofísica
Mexicana, GEOS, v. 21, p. 206.
Carranza-Castañeda, Oscar, and Miller, E.W., 1996, Hemphillian and
Blancan Felids from Central Mexico: Journal of Paleontology,
v. 70, no. 3, p. 509-518.
Carrasco-Nuñez, G.J.; Milán, M.; and Verma, S.P., 1989, Geología del
volcán El Zamorano, Estado de Querétaro: Universidad
Nacional Autónoma de México, Instituto de Geología, Revista,
v. 8, p. 194-201.
Cerca-Martínez, M.; Aguirre-Díaz, G.J.; and López-Martínez, M.,
2000, The geologic evolution of southern Sierra de Guanajuato,
Mexico-A documented example of the transition from the
Sierra Madre Occidental to the Mexican Volcanic Belt:
International Geology Review, v. 12, no. 2, p. 131-151.
Chiodi, M.; Monod, O.; Busnardo, R.; Gaspard, D.; Sánchez, A.; and
Yta, M., 1988, Une discordance anté-albienne datée par una
faune d’ammonites et de brachiopodes de type téthysien au
Mexique central: Geobios, no. 21, p. 125-135.
Corona-Chávez, P., 1988, Análisis estratigráfico-estructural de la porción centro-oriental de la Sierra de Guanajuato: México, D.F.,
Instituto Politécnico Nacional, Escuela Superior de Ingeniería
y Arquitectura, BS thesis, 60 p. (unpublished).
145
Dávila-Alcocer, V.M., and Martínez-Reyes, Juventino, 1987, Una
edad cretácica para las rocas basales de la Sierra de
Guanajuato: Universidad Nacional Autónoma de México,
Instituto de Geología, Simposio sobre la geología de la Sierra
de Guanajuato, resúmenes, p.19-20 (abstract).
Demant, Alain, 1978, Características del Eje Neovolcánico
Tansmexicano: Universidad Nacional Autónoma de México,
Instituto de Geología, Revista, v. 2, p. 172-187.
Echegoyén, J.; Romero-Martínez, S.; and Velázquez-Silva, S., 1970,
Geología y yacimientos minerales de la parte central del distrito minero de Guanajuato: Boletín Consejo de Recursos
Minerales no Renovables, v. 75, p. 35.
Edwards, D.J., 1955, Studies of some early Tertiary red conglomerates
of central Mexico: U.S. Geological Survey, Professional Paper
264-H, p. 153–185.
Ferrari, Luca; López-Martínez, M.; Aguirre-Díaz, G.J.; and CarrascoNúñez, G., 1999, Space-time patterns of Cenozoic arc volcanism in central Mexico, from the Sierra Madre Occidental to the
Mexican Volcanic Belt: Geology, v. 27, p. 303–306.
Freundt, A.; Wilson, C.J.N.; and Carey, S.N., 2000, Ignimbrites and
block-and-ash flow deposits, in Sigurdsson, H., Encyclopedia
of Volcanology: Academic Press, p. 581–599.
Gross, W.H., 1975, New ore discovery and source of silver-gold veins,
Guanajuato, Mexico: Economic Geology, v. 70, p. 1175–1189.
Henry, C.D., and Aranda-Gómez, J.J., 1992, The real southern Basin
and Range; mid- to late Cenozoic extension in Mexico:
Geology, v. 20, p. 701–704.
Henry, C.D., and Aranda-Gómez, J.J., 2000, Plate interactions control
middle-late Miocene, proto-Gulf and Basin and Range extension in the southern Basin and Range: Tectonophysics, v. 318,
p. 1–26.
Kowallis, B.J.; Swisher, C.C.III; Carranza-Castañeda, Oscar; Miller,
W.E.; and Tingey, D.G., 1998, Fission-track and single-crystal
39Ar-40Ar laser-fusion ages from volcanic ash layers in fossilbearing Pliocene sediments in central Mexico: Revista
Mexicana de Ciencias Geológicas, v. 15, no. 2, p. 157–160.
Labarthe-Hernández, Guillermo; Tristán, M.; and Aranda-Gómez,
J.J., 1982, Revisión estratigráfica del Cenozoico de la parte
central del Estado de San Luis Potosí: Instituto de Geología y
Metalurgia, Folleto Técnico, v. 85, 208 p.
Lapierre, H.; Ortiz, E.; Abouchami, W.; Monod, O.; Coulon, C.; and
Zimmermann, J.L., 1992, A crustal section of an intra-oceanic
island arc; the Late Jurassic-Early Cretaceous Guanajuato magmatic sequence, central Mexico: Earth and Planetary Science
Letters, v. 108, p. 61-77.
Martínez-Reyes, Juventino, 1992, Mapa geológico de la Sierra de
Guanajuato con resumen de la geología de la Sierra de
Guanajuato: Universidad Nacional Autónoma de México,
Instituto de Geología, Cartas geológicas y mineras 8, escala
1:100,000.
Martínez-Reyes, Juventino, and Nieto-Samaniego, A.F., 1992, Efectos
geológicos de la tectónica reciente en la parte central de
México: Universidad Nacional Autónoma de México, Instituto
de Geología, Revista, v. 9, p. 33–50.
Monod, O.; Lapierre, H.; Chiodi, M.; Martínez, Juventino; Calvet, P.;
Ortiz, E.; and Zimmermann, J.L., 1990, Reconstitution d’un
arc insulaire intraocéanique au Mexique central; la séquence
volcano-plutonique de Guanajuato (Crétacé Inférieur):
Comptes Rendus de l’Académie des Sciences de Paris, v. 310,
p. 45–51.
146
Múgica, R., and Albarrán, J., 1983, Estudio petrogenético de las rocas
ígneas y metamórficas del Altiplano: Mexico, Instituto
Mexicano del Petróleo, Informe del proyecto C-1156, 78 p.
(unpublished).
Nieto-Samaniego, A.F., 1990, Fallamiento y estratigrafía cenozoicos
en la parte sudoriental de la Sierra de Guanajuato: Universidad
v. 9, p. 146-155.
Nieto-Samaniego, A.F.; Macías-Romo, Consuelo; and AlanizAlvarez, S.A., 1996, Nuevas edades isotópicas de la cubierta
volcánica cenozoica de la parte meridional de la Mesa Central,
México: Revista Mexicana de Ciencias Geológicas, v. 13, no.
1, p. 117–122.
Ortega-Gutiérrez, Fernando; Mitre-Salazar, L.M.; Roldán-Quintana,
Jaime; Aranda-Gómez, J.J.; Morán-Zenteno, D.J.; AlanizÁlvarez, S.A.; and Nieto-Samaniego, Á.F., 1992, Carta geológica de la República Mexicana: México, Universidad Nacional
Autónoma de México, Instituto de Geología, Secretaría de
Energía, Minas e Industria Paraestatal, Consejo de Recursos
Minerales, 1 sheet.
Ortiz-Hernández, L.E.; Chiodi, M.; Lapierre, H.; Monod, O.; and
Calvet, Ph., 1990 (1992), El arco intraoceánico alóctono
(Cretácico Inferior) de Guanajuato-Características petrográficas, geoquímicas, estructurales e isotópicas del complejo filoniano y de las lavas basálticas asociadas; implicaciones geodinámicas: Universidad Nacional Autónoma de México,
Instituto de Geología, Revista, v. 9, no. 2, p. 125–145.
Ortiz-Hernández, L.E., and Martínez-Reyes, Juventino, 1993,
Geological structure, petrological and geochemical constraints
for the centralmost segment of the Guerrero Terrane (Sierra de
Guanajuato, central Mexico): Guidebook of field trip C, First
Circum-Pacific and Circum-Atlantic Terrane Conference,
Guanajuato (México), November 5–22, 25 p.
Ortiz-Hernández, L.E.; Flores-Castro, K.; and Acevedo-Sandoval,
O.A., 2002, Petrographic and geochemical caracteristics or
upper Aptian calc-alkaline volcanism in San Miguel de Allende
(Guanajuato state), Mexico: Revista Mexicana de Ciencias
Geológicas, v. 19, p. 87–91.
Quintero-Legorreta, Odranoel, 1992, Geología de la región de
Comanja, estados de Guanajuato y Jalisco: Universidad
v. 10, no. 1, p. 6–25.
Pasquaré, G.; Forcella, F.; Tibaldi, A.; Vezzoli, L.; and Zanchi, A.,
1986, Structural behavior of a continental volcanic arc; the
Mexican Volcanic Belt, in Wezel, F.C. ed., The origin of arcs,
Developments in Geotectonics, Elsevier, p. 509–527.
Pasquaré, G.; Ferrari, Luca; Perazzoli, V.; Tiberi, M.; and Turchettti,
F., 1987a, Morphological and structural analysis of the central
sector of the Transmexican Volcanic Belt: Geofísica
Internacional (Mexico), v. 26, p. 177–194.
Pasquaré, G.; Vessoli, L.; and Zanchi, A., 1987b, Morphological and
structural model of the Mexican Volcanic Belt: Geofísica
Internacional (Mexico), v. 26, p. 159–176.
Pasquaré, G.; Garduño, V.H.; Tibaldi, A.; and Ferrari, Luca, 1988,
Stress pattern evolution in the central sector of the Mexican
Volcanic Belt: Tectonophysics, v. 146, p. 353–364.
Pérez-Venzor, J.A., 1997, Estudio de la evolución geológica del complejo volcánico Palo Huérfano, Mpio. de San Miguel Allende,
Gto. México, D.F., Universidad Nacional Autónoma de
México, MSc thesis, 95p. (unpublished).
Pérez-Venzor, J.A.; Aranda-Gómez, J.J.; McDowell, F.W.; and
Solorio-Munguía, J.G., 1996, Geología del Volcán Palo
Huérfano, México: Revista Mexicana de Ciencias Geológicas,
v. 13, p. 174–183.
Randall-Roberts, J.A.; Saldaña-A., E.; and Clark, K.F., 1994,
Exploration in a volcano-plutonic center at Guanajuato,
Mexico: Economic Geology, v. 89, p. 1722–1751.
Tristán-González, Margarito, 1986, Estratigrafía y tectónica del
graben de Villa de Reyes, en los estados de San Luis Potosí y
Guanajuato: Universidad Autónoma de San Luis Potosí,
Instituto de Geología, Folleto Técnico, v. 107, p. 91p.
Sedlock, R.L.; Ortega-Gutiérrez, Fernando; and Speed, R.C., 1993,
Tectonostratigraphic terranes and tectonic evolution of Mexico:
Geological Society of America, Special Paper, v. 278, 153 p.
Stein, G.; Lapierre, H.; Monod, O.; Zimmermann, J.L.; and Vidal, R.,
1993, Petrology of some Mexican Mesozoic plutons-sources
and tectonic environments: Journal of South American Earth
Sciences, v. 7, no. 1, p. 1–7.
Tardy, Marc; Lapierre, H.; Boudier, J-L.; Yta, M.; and Coulon, Ch.,
1991, The Late Jurassic-Eearly Cretaceous arc of western
Mexico (Guerrero terrane); origin and geodynamic evolution:
Universidad Nacional Autónoma de México, Instituto de
Geología, Convención sobre la evolución geológica de
México y I Congreso mexicano de Mineralogía, Memoria, p.
213–215.
Tardy, Marc, Lapierre, H., et al., 1994, The Guerrero suspect terrane
(western Mexico) and coeval arc terranes (the Greater Antilles
and the Eastern Cordillera of Colombia); a late Mesozoic intraoceanic arc accreted to cratonal America during the
Cretaceous: Tectonophysics, v. 230, p. 49–73.
Valdez-Moreno, G.; Aguirre-Díaz, G.J.; and López-Martínez, M.,
1998, El volcán La Joya, estados de Querétaro y GuanajuatoUn estratovolcán miocénico del Cinturón Volcánico Mexicano:
Revista Mexicana de Ciencias Geológicas, v. 15, p. 181–197.
Zimmermann, J.L.; Stein, G.; Lapierre, H.; Vidal, R.; Campa, M.F.;
and Monod, O., 1990, Données géochronologiques nouvelles
sur les granites laramiens du centro et l’ouest du Mexique
(Guerrero et Guanajuato): Société Géologique de France,
Réunion des Sciences de la Terre, 13, Grenoble, France, p. 127
(abstract).
D AY 1 : A N O V E RV I E W O F T H E G E O L O G Y
T H E G U A N A J U AT O M I N I N G D I S T R I C T
OF
MENT: THE
JURASSIC-CRETACEOUS VOLCANIC ARC
(FORE-ARC) SEDIMENTARY BASIN
ADJACENT
14Q0266176; 2328677)
147
AND ITS
(UTM
Refer to Figures 5 and 6 for locations of stops.
Km 0.0: We begin at the Hotel Parador San Javier.
Heading north on the paved road toward Dolores Hidalgo
(Federal Highway 110), the highway passes through the
Tertiary redbeds of the Guanajuato Conglomerate. Km 1.9: A short distance before the town of La
Valenciana, looking toward the east we can see a
panoramic view of an outcrop of the Veta Madre faultvein. The structure was mined at the surface and the old
works resemble triangular facets. Closer to the road, in
the stream, are the tailing ponds of the Cooperativa Santa
Fe de Guanajuato.
Km 2.4: La Valenciana town. To the right is an imposing
sixteenth century church, which was built on the trace of
the NE-trending Aldana Fault, which puts into contact
the Eocene redbeds of the Guanajuato Conglomerate and
the Mesozoic Formations (Figure 5). Once we pass the
town, the road continues uphill through a diabasic dike
complex emplaced in dioritic to tonalitic host rocks
(shown in Figure 5 as La Palma Diorite), part of the
Mesozoic volcanoplutonic sequence. The outcrops of the
complex are located west of the road.
Km 3.2: Near the Camino de Guanajuato Hotel, the road
crosses the trace of the Veta Madre fault, which puts into
contact rocks of the volcanoplutonic sequence of the
Guanajuato Arc and pelagic sediments of the Arperos
Basin. From this point on, the highway was built on
intensely deformed rocks of the Arperos Basin, mostly
slates and thin-bedded limestones.
Km 3.5: To the left is the road that leads to the La
Esperanza Dam and the Los Insurgentes (DIF) campground. As we go uphill on Highway 110, a panoramic
view of the city of Guanajuato and the town of La
Valenciana can be seen to the right.
Km 4.45: To the left, there is a small flat area where we
will pull off the road and have our first stop. From this
locality we can obtain a broad overview of the most relevant rock units and Cenozoic structures exposed in the
region.
STOP 1-1. INTRODUCTION TO THE GEOLOGY OF THE
GUANAJUATO MINING DISTRICT, PANORAMIC VIEW OF THE
CITY, LITHOLOGIES AND CONTACT RELATIONSHIPS IN THE BASE-
The most important regional features seen in this
overview are as follows:
—To the south, the city of Guanajuato was built in a
basin occupied by the redbeds of the Guanajuato
Conglomerate. The southeastern part of the city is
flanked by near-vertical cliffs where the lowermost
Tertiary volcanic formations overlie the Guanajuato
Conglomerate (Figs. 5, 7). Behind those mountains is the
wide valley known as El Bajío. The depression in which
Guanajuato is limited to the west by the NE-trending
Aldana Fault and to the north by the NW-trending Veta
Madre fault (Figure 6). Southeast of Stop 1-1 are the hills
Cerro Chichíndaro and Cerro de Sirena, with their northeastern flanks displaced by the Veta Madre fault, which
in that place marks the limit between the Mesozoic and
Cenozoic units. From this point it is also worthwhile to
observe the marked changes in thickness of some of the
mid-Tertiary volcanic units, such as the Bufa Ignimbrite,
which in the vertical cliffs south of Stop 1-3 exceeds 300
meters and in Cerro de Sirena is ≤10 meters. We believe
this dramatic northward decrease in thickness indicates
active erosion and/or caldera-wall faulting at the time of
Bufa volcanism, or possibly deposition of the Bufa ashflow tuff on a surface made very irregular by regional
faulting.
—To the east, we see the mountain range known as the
Sierra de Santa Rosa, and it is partially covered by
Tertiary volcanic units which dip gently (~12 degrees)
toward the northeast. Compared with the sharp southwestern limit of the Sierra, determined by the El Bajío
fault, the northeastern boundary is poorly defined, and
the volcanic rocks gradually merge with the thick
deposits of gravel that partially fill the Río Laja Basin.
The overall geomorphologic picture of the region suggests that the Sierra is located at the southern end of a
large block tilted to the northeast during the late
Cenozoic.
—To the north, in the background are exposures of the
Cerro Pelón Tonalite (Figure 5), which stands out as the
white ground without vegetation. In the foreground is
Cerro El Plomo, which is made up of a terrigenous flysch
sequence; between that hill and us, in the bottom of the
stream valley, is La Esperanza Dam.
148
—To the west, in the background, is Cerro El Cubilete,
crowned by a shrine topped with a large statue of Christ.
The church was built on top of Late Cenozoic andesite
and gravel, which in turn overlie the Mesozoic volcanoplutonic rocks of the Guanajuato Arc. Between El
Cubilete and Stop 1-1 is an unforested hilly country with
extensive outcrops of the dike complex (shown in Figure
5 as La Palma Diorite) and metalavas and metatuffs
which form parts of both the volcanoplutonic association.
In that area there are also extensive outcrops of the volcanosedimentary association (Figure 5). Description of the outcrops at Stop 1-1
In the road cuts there are examples of the complex contact relationships among the various volcanic components of the Mesozoic basement, common in all parts of
the Sierra de Guanajuato. They also show some of the
characteristic lithologies of the sedimentary basin near
the volcanic arc (Tardy and others, 1991). The outcrop
includes interbedded pelitic and carbonate sediments of
the Esperanza Formation (Echegoyén, 1970), submarine
andesite flows and tuffs, and tonalitic intrusives. In this
outcrop the calcareous-to-argillaceous sedimentary rocks
are intensely deformed, displaying a well-developed
schistosity. In the argillaceous rocks, crenulations striking approximately N-S, associated with a pre-Laramide
deformation, are observed. During Laramide deformation these structures were compressed, producing microfolds which are clearly observed along the outcrop. The
axial planes of these microfolds have an average strike of
N35-40W and an average dip to the northeast of less than
45 degrees. Other prominent structural features include
incipient shear zones in the cores of some of the larger
folds, and lenses or boudins of weakly metamorphosed
marls dispersed among the more phyllitic argillaceous
layers. Another notable feature is the presence of nearvertical extension fractures. These may be associated
with mid-Tertiary Basin-Range extension, or possibly
with a stress field related to the accretion of the Arperos
forearc basin sequence to the Guanajuato Arc during final
assembly of this portion of the Guerrero Terrane, just
prior to its obduction onto the edge of North America. Volcanic rocks metamorphosed to greenschist
facies are found both in tectonic contact and in depositional contact with the calcareous metasediments. The
protoliths of these rocks probably were principally
andesitic lava flows; however, metatuffs are also abundant. The metatuffs form sequences of thin layers in
which basal crystal concentrations are easily recognizable. Some layers also have a marked gradation in particle size from lapilli-rich bases to ash-rich tops. At this
site it is possible to observe one of the many thrust contacts in the Mesozoic sequence. We speculate that it was
folded, probably during the Laramide Orogeny. Along
this thrust an allochthonous slice of intrusive rock, apparently a leucotonalite, which now has a mylonitic foliation, was transported over rocks of the volcanosedimentary association described previously. Several Tertiary
normal faults displace the thrust; in some places adjacent
to these later faults, the thrust and the rocks of the road
cut in general are drag-folded into nearly vertical positions.
The basement-rock lithologies observed in this
roadcut will be easily identifiable as clasts in the
Guanajuato Red Conglomerate and also as accidental
lithics in the volcanic rocks of the Mining District.
We will retrace our route from this stop, returning to
the city of Guanajuato. Before arriving at our base hotel,
we will take the Panorámica Highway heading southwest
until we reach the dirt road near the School of Mines of
the University of Guanajuato.
STOP 1-2: GUANAJUATO RED CONGLOMERATE
AND
~49 MA
LAVAS EXPOSED IN A CUT ON A DIRT ROAD LOCATED CLOSE TO
THE
SCHOOL OF MINES (UTM 14Q0264773;2326644)
At this site are outcrops of greenish andesitic lavas intercalated with the Guanajuato Red Conglomerate (GRC).
The conglomerate, with attitude N4W, 30SW, shows
noticeable variations in grain size, sorting, thickness of
individual strata, and nature of the clasts. Near the
Panorámica is an alternation of conglomeratic coarsegrained sandstones with fissile shales in strata from10 to
30 centimeters in thickness. As we move to the west, the
grain size increases considerably and the thickness of the
layers goes up to about a meter, and in some layers normal grading and/or the presence of paleochannels is
apparent.
In this outcrop one also observes at least four lava
flows, each with an autobrecciated base and a vesicular
top. The flows vary from the massive type to the blocky
type. The rock is generally fine-grained, with scarce phenocrysts of plagioclase and altered mafic minerals. At the
base of the sequence of flows we can see a tuff a little less
than a meter thick, composed principally of coarse to fine
ash and oxidized mafic minerals. The tuff is concordant
with the underlying GRC. The lavas are completely
propylitized, turning them into green rocks very similar
in aspect to the basement metalavas that we saw at Stop
1-1. Aranda-Gómez and McDowell (1998) obtained a KAr age (whole-rock) of ~49 Ma on one of the outcrops of
andesite intercalated in the GRC. This age, together with
the paleontological ages reported by Edwards (1955) and
the K-Ar age of the Bufa Ignimbrite (~37 Ma; Gross,
1975), permits us to know the age of deposition of the
GRC, and therefore to have an idea of the time at which
the Guanajuato Basin formed (Middle Eocene to Early
Oligocene). The GRC, which overlies these lavas, is rich
in large boulders of andesite, similar to the underlying
andesite or to the metalavas of the basal complex. It also
contains clasts of felsic volcanic rocks of unknown
provenance and some fragments of limestone.
STOP 1-3: SUBSTATION CFE AND LA CUEVA. CONTACT
THE GRC WITH THE LOSERO FORMATION AND THE BASE
THE BUFA IGNIMBRITE (UTM 14Q0266028; 2323640)
OF
OF
We will return to the base hotel and from there we will
take the road toward the center of the city. We will cross
Guanajuato heading toward the Silao exit. Before arriving at the exit we will take the overpass to Pozuelos, in
order to get onto the Panorámica in its southern part. We
pass the ISSSTE Clinic and stop at the CFE electrical
substation. From this saddle (UTM 14Q0266028; 2323640),
we can take in a panoramic view of Cerro de La Bufa, the
place from which the name of the Bufa Ignimbrite comes,
on one side (south). On the other side (north), we can see
the city of Guanajuato from a point opposite to that of
Stop 1-1. From this point we see the contact between the
GRC and the overlying units, the Losero Formation and
the Bufa Ignimbrite. The contact is very clearly exposed,
given the scarce vegetation and the marked color contrast
between the units, dark red for the GRC and light green
and yellowish green for the Losero-Bufa package. From
here we can also observe the change in dip of the upper
layers of the GRC with respect to the lower layers. The
dip becomes gradually gentler upward, and at the top is
even almost concordant with the subhorizontal layers
(~16 degrees) of the Losero Formation. This gradual
149
change in the dip is interpreted as a rollover fold in the
Tertiary sequence by Aranda-Gómez and McDowell
(1998). These authors argue that the GRC and the volcanic sequence were accumulated at the same time that
intense normal faulting was going on in the region.
From here we will go up to the contact area of the
three units (UTM 14Q0266471; 2323485). Then, we will
walk along the contact between the GRC and the Losero
to a place known as La Cueva (UTM 14Q0265958;
23230009), that is a somewhat tabular excavation made
in order to extract sheets and blocks of the Losero
Formation. Losero by reason of its characteristic green
color and finely laminated layers with graceful dune bedforms has been used as an ornamental stone in constructions around the region. The quarrying of the Losero is no
longer going on at La Cueva, and now it is a small chapel
traditionally visited during the Holy Week holidays.
Owing to the workings of this quarry, the Losero is
unusually well exposed, and it is possible to observe
details of the sequence of surge deposits in the Losero on
mutually perpendicular surfaces, as well as the contact
between the Losero surge beds and the Bufa Ignimbrite.
The GRC-Losero Transition: At this locality (UTM
14Q0266157; 2323097) we have a well-exposed section
which shows the transition from the GRC to the Losero
Formation and the contact between the Losero Formation
and the Bufa Ignimbrite. The upper layers of the GRC are
strata of fine gravel and coarse sands with rather thin
bedding (10-20 centimeters) which give way upward to
red siltstone. The sequence continues with a mixed interval of green and red siltstone with cross-bedding, which
changes in an almost imperceptible way to the entirely
green layers of the Losero. Therefore the contact between
the GRC and the Losero is transitional and concordant.
Characteristics and origin of the Losero: Along the outcrops on this road, the Losero has a thickness that is variable, from 0 to 10 meters, and it consists principally of a
rhythmic sequence of surge layers with intercalations of
epiclastic deposits of well-sorted sands and silts. These
deposits form fine laminations with thicknesses from a
few millimeters up to 3 centimeters. The surge layers
have both dune and antidune forms (very low -amplitude
and long-wavelength dunes with well-developed stossside accretion and lee-side erosion of material). These
features mostly indicate transport from SE to NW,
150
assuming that the temperature of the pyroclastic current
from which they were deposited was above the boiling
point of water, so that the crests of successive dunes
upward in each duneform retreat toward the source. The
surge deposits are composed of pyroclastic material that
varies from very fine ash to lapilli. The uppermost part of
the Losero includes some deposits of very fine ash in
which normal grading can be seen. We think that many of
the pyroclastic deposits described above accumulated in
water, possibly in ephemeral shallow lakes. This interpretation is based on the presence of layers of well-sorted sandstone and siltstone, and because we also occasionally observe traces of surge features in some of these
strata. This interpretation is congruent with the lithology
of the uppermost portion of the GRC; however, it is also
possible that these apparently detrital layers are in fact
the planar facies corresponding to the sandwave facies
described above.
The Losero-Bufa Contact: The Bufa Ignimbrite in many
places seems to overlie the Losero concordantly, but in
others it is easy to see layers of the Losero truncated by
the base of the ignimbrite. One can also see surfaces with
raindrop pits in the upper layers of the Losero, which
implies that there was a time lapse between the activity
that produced the Losero and that which produced the
Bufa Ignimbrite. On the other hand, there are no
reworked deposits or paleosoils between the two units.
Thus we infer that the Bufa Ignimbrite was emplaced
shortly after the accumulation of the Losero surge layers.
In fact, it is possible that the Losero and the Bufa are both
products of the same source and that the Losero represents the initial phases of the paroxysmal eruption that
produced the Bufa. The contact forms a planar surface
with an average attitude of N30W, 18NE. The lack of
soft-sediment load structures in the Losero suggests that
it already had considerable bearing strength at the time
that the great weight of the Bufa was deposited on top of
it.
Characteristics of the lower portion of the Bufa
Ignimbrite: In these outcrops at La Cueva the Bufa is
approximately 400 meters thick. We will only be able to
study the lowermost part of the unit. The features displayed by the Bufa in this section are those of an ignimbrite with a great deal of kinetic energy, but little thermal energy, since the degree of welding is in general fair-
ly low. In the lower part of the Bufa, just above the contact with the Losero, is a zone approximately 2 meters
thick which is rich in lithics and has flattened pumices or
green fiamme. It changes gradually upward to a partially
welded, lithic-poor zone several meters thick. This zone
is overlain by a zone with abundant hollow pits, which
range in size from golf balls to soccer balls. These pits
result from differential erosion of pumice clasts with
respect to matrix, the pumices in this case being more
easily eroded. Above the pitted layer, the ignimbrite
changes to a highly silicified zone gray in color with
black spots and patches of iron oxide. Silicification is
pervasive in this zone, transforming it into a highly erosion-resistant rock which projects out over the lower, less
silicified part making a prominent overhang. The original
texture of the silicified zone was totally obliterated, and
secondary quartz is abundant. As primary minerals the
ignimbrite contains fairly abundant euhedral biotite, sanidine and quartz.
As was mentioned at Stop 1-1, the Bufa Ignimbrite
has great lateral variations in thickness within the
District, and outside of the District it seems to be absent.
Within a horizontal distance of less than 5 kilometers its
thickness varies from a maximum of 400 meters here at
Las Cuevas to a minimum of 0–10 meters at Cerro de
Sirena. It is possible that one or more of the curvilinear
faults that separate this outcrop from Cerro de Sirena (the
Amparo and San Clemente faults and the northeastern
branch of the Veta Madre) form part of the northern margin of a caldera associated with the Bufa eruption,
although the source(s) of the Bufa are not known. Ash
flows exposed near the intersection of the Aldana and El
Bajío faults are thought to correlate with the Cuatralba
Ignimbrite, dated at 30 Ma, and not with the Bufa.
According to the work of Martínez-Reyes (1992), the
Cuatralba Ignimbrite is widely distributed west of the
Villa de Reyes graben, as well as along the downthrown
block of the El Bajío fault (including two small outcrops
within the La Sauceda graben, SSE of La Cueva).
Likewise the Calderones Formation, which has a very
distinctive lithology, has not been observed in areas outside the block bounded by the Rodeo-Yerbabuena fault,
the La Sauceda graben, and the La Sierra vein system.
We will return to the vehicles and take the Panorámica
Highway toward the La Olla Dam (to the east). This construction dates from the 1700’s, and for many years it was
the only source of water for the city. It also played an
important role in controlling floods (the city was
destroyed several times by flooding). Upon arriving at
the dam, we will continue a bit further toward the monument to Miguel Hidalgo and the square at the San
Renovato Dam.
Km 0.0: We will take the cobblestone road (which later
turns into a dirt road) toward El Cubo. Km 1.0: Contact between the GRC and the Losero
Formation. A few meters up the road the Bufa Ignimbrite
comes into view. These contacts are exposed on both
sides of the road and in the cuts on the other side of the
San Renovato Dam. As we go further up the road we can
look down into the canyon on the right. At the bottom of
this canyon you can see the green layers of the Losero
under the Bufa Ignimbrite. The road crosses a zone where
the Bufa has an interval of crude columnar jointing. Here,
as everywhere, the Bufa has been eroded to form spectacular cliffs.
Km 1.4: The road crosses the Arroyo de Los Rieles. We
will park in a spot in which the road widens next to a tiny
chapel.
STOP 1-4: CONTACT BETWEEN THE UPPER PART OF THE BUFA
AND THE CALDERONES FORMATION. DISTAL FACIES OF THE
CALDERONES SEQUENCE IN THE ARROYO DE LOS RIELES
(UTM 14Q0268222; 2324110)
We will walk upstream in the arroyo for a short distance
in order to be able to observe the uppermost part of the
Bufa Ignimbrite and the lower part of the Calderones
pyroclastic sequence. The Bufa Ignimbrite here at the top of the unit is
pink (color intensifies upward in the uppermost part) and
very fine-grained. It is weakly welded, but a bit indurated due to silicification. Some white pumice clasts can be
seen, but both pumicees and lithics (red aphyric rhyolite)
are scarce. The contact between the Bufa and the
Calderones is poorly exposed, but we will see it later in
this same stop, in a large roadcut uphill from the curve
where the vehicles are parked. The base of the Calderones (UTM 14Q0268369;
2324216) is a well-stratified deposit, with individual layers ranging from 5 to 30 centimeters in thickness. The
bottom 3 meters are characterized by relatively thick bedding (up to 0.3 meters) and by the presence of abundant
151
angular fragments of pale reddish purple dacite. These
distinctive clasts were probably derived from the domes
of the Peregrina Dome Field, which we will visit in the
last stop (1-7) of the day. Other recognizable clasts
include those from the GRC and some andesite chips.
The high content of lithics and their angularity impart a
roughness to exposed surfaces in this part of the
Calderones sequence. We interpret this part of the
sequence as the distal deposits of several thin pyroclastic
flows. Above the clast-rich basal beds, the unit is composed of generally finer-grained and more finely laminated green layers (less than 30 centimeters thick, on
average), with prominent cross-bedding. In these layers
dune forms with stoss-side accretion of laminae are clearly exposed. Some of the dunes have pebble trains, and
dune-regression patterns are compatible with a NE-toSW transport direction. We interpret these beds as surge
layers; they are interbedded with layers that lack crossbedding but are graded with respect to clast size, which
we interpret as probable fall deposits. The colors of individual layers in this interval include both purplish red and
grayish green, with green predominating. One exceptionally fine-grained, planar-bedded interval about a meter
thick may be a planar-facies surge deposit or a fall
deposit composed of fine ash; its orientation is N45W,
10NE. Further upstream are more pyroclastic flow
deposits with cross-bedding, and some thicker (0.5
meters to several meters) ignimbrites. One such unit contains a feature typical of the Calderones distal facies, a
wedge of massive tuff occupying a paleochannel cut into
stratified deposits. At this point in the arroyo, the packages of ignimbrites form high ledges (3-5 meters) which
make it difficult to climb further upstream. The entire
Calderones section at this point is bright green, the characteristic color imparted to the unit by pervasive chloritization of all original glass. We will return to the El Cubo road and walk up it
for a couple of hundred meters in order to study a cut,
which has a good exposure of the contact between Bufa
and Calderones. At UTM 14Q0268457; 2324077 we can see that the
pyroclastic flows at the base of the Calderones filled a
broad paleochannel developed in the upper part of the
Bufa Ignimbrite. Therefore the contact between these
units is locally an erosional disconformity. Unlike most
channel fillings, this one does not have any of the typical
deposits, such as gravels and sands, nor is there any evi-
152
dence of a paleosoil on top of Bufa. The Calderones pyroclastic flow deposits rest directly on the Bufa Ignimbrite;
they consist of a series of thin relatively coarse-grained
layers, slightly wavy, with cross-bedding and with lithics
from the Peregrina Dome and the GRC, similar to the less
well-exposed basal beds in the arroyo. This initial
sequence, which is seen here thinning and becoming
finer-grained toward the channel margins, probably
resulted from the initial blasts and/or surges related to the
channel-filling pyroclastic flow(s) which overlie it. About
midway between the bottom and the top of the roadcut,
there is a layer of large clasts at the base of a thicker (more
than 5 meters) pyroclastic flow. In total the Calderones
sequence exposed in this cut is approximately 12 meters
thick in the central part of the channel filling. Km 3.9: Junction with the road to the Las Torres Mine.
This mine is one of the most recent discoveries (Gross,
1975) along the Veta Madre. It is also one of the largest
and most modern operations in the District. STOP 1-5: CONTACT BETWEEN THE CALDERONES PYROCLASTIC
ROCKS AND THE CEDRO ANDESITE (UTM 14Q0268833;
2323324)
About 100 meters west of the bus stop, the contact
between the two units is exposed. In an interval of about
12 meters, it is possible to see that at the top of the
Calderones is a sequence (~3 meters thick) of thin layers
composed of relatively crystal-rich tuffs with highly
vesiculated and intensely palagonitized glassy matrix
material. These yellowish brown layers are interstratified
with the more typical fine-grained green layers in the
lower part of the transitional interval. Just at the base of
the Cedro lava flows is a horizon of very thin layers with
well-formed dessication cracks. Above these layers rests
an andesite flow with well-developed spheroidal weathering and small pillows (?), which changes gradually
upward into massive andesite. We interpret the observed
features at the flow base as evidence that it interacted
with water. In a small quarry located to the north of the
bus stop it is possible to see that the first andesite flow at
the base of the Cedro is overlain by another flow with
characteristics similar to those described above.
Km 4.5: Looking eastward from this point you can see
the village of El Cedro in the bottom of the canyon; the
name of the Cedro Andesite was taken from this location.
In the same direction, on the other side of the valley, is
the Sierra de La Leona (i.e. the hill with the cross on top
at the northern end of a long ridge). The ridge is composed of red layers of the GRC at its base and of Losero
Formation and Bufa Ignimbrite along the summit. The
Cenozoic sequence is repeated by the Veta Madre fault,
which here has considerable throw (hundreds of meters),
bringing the Cedro Andesite down against the GRC (Stop
1-6).
Km 4.7: Immediately west of the road there is an
embankment of material excavated in the Cedro
Andesite. The material quarried from here is used to contain the tailings ponds behind the adjacent dam.
Km 5.0: Here we will park at the side of the road to
inspect the Veta Madre fault zone.
STOP 1-6: THE VETA MADRE FAULT AND THE RELATIONSHIP
BETWEEN CALDERONES AND CEDRO FORMATIONS (UTM
14Q026497; 2324366)
The Veta Madre crops out here along the road cut. In this
region the Veta Madre branches, surrounding the ridge
capped by the Chichíndaro Rhyolite. In the road cut the
fault has an approximately E-W orientation, but on the
regional scale it strikes NW-SE. Walking along the roadcut, you can see tectonic contacts between the GRC and
Calderones, and Calderones with Cedro, even though the
rocks in this outcrop are intensely altered and brecciated.
Km 5.8: North of this point we have Cerro Chichíndaro,
which is crowned by a dome or domes of rhyolitic lava
with a K-Ar age of ~32 Ma (Gross, 1975), and it is the
type locality of the Chichíndaro Rhyolite. The age of the
Veta Madre fault is bracketed between the age of the
Chichíndaro Rhyolite and the age of the mineralization
along it (K-Ar, ~29-27 Ma; Gross, 1975).
Km 5.7: We cross the pass between Cerro Chichíndaro
and the hill to the east. The Las Escobas fault crosses this
pass; this structure puts the basal Mesozoic complex in
tectonic contact with the redbeds of the GRC.
Km 6.6: Intersection of road to El Cubo with road north
to Peregrina. We turn left onto the road that goes to the
Peregrina Mine. Km 9.2: We will stop here, pulling off the road near the
top of the rise.
153
STOP 1-7: THE PEREGRINA DOME COMPLEX AND ITS RELATIONSHIP TO THE CALDERONES SEQUENCE (UTM
14Q0271496; 2326250)
We will go out of the base hotel headed for the La Olla
Dam and will take the road to El Cubo. Refer to Figure 6
for locations of stops.
From this vantage point we can see, first, a wicked good
panoramic view of the Peregrina Dam and the northwestern part of the Peregrina Dome Complex. Second, outcrops of the dacitic to latitic phase of the Peregrina and of
a small block-and-ash flow derived from the collapse of
one of the domes. Study of these outcrops reveals a
dynamic scenario of growth and destruction of small
domes, repeatedly, which produced several block-andash flows that filled paleochannels and/or a paleograben
just to the south of these domes. The contact between
these two units, Peregrina domes and Calderones tuffbreccias and other proximal deposits, seems to be a complex and repetitive one, as the pyroclastic flows produced
by the Peregrina domes are intercalated with, and indeed
form part of, the Calderones sequence. South of here,
many of the flow units observed in the Calderones are
rich in angular lithic fragments of this dacitic phase of the
Peregrina, as we will see in the Arroyo de Los Silvestres,
later in the trip (Stop 2-4).
Looking toward the Peregrina Dam, you can see
two small hills on the skyline. In the sides and tops of
these hills, flow foliation typical of domes can be seen;
that is to say, there is a concentric pattern of pseudostratification that gradually becomes steeper from the exterior to the interior of the dome. This flow-banding is nearly vertical around the Peregrina Dam.
Where we are standing, we can see in detail the contact relationships between the external part of one of the
domes and the associated pyroclastic deposits. We can see
that the rock making up the internal part of the dome is composed of a pale grey porphyritic dacite, with well-developed
flow-banding. It has phenocrysts of plagioclase and quartz
in a devitrified groundmass. The outer part of the dome consists of a totally devitrified carapace made up of the brecciated equivalent of the grey porphyry in the interior. A few
meters beyond the outcrop of the intact dome rocks, the
proximal facies of the dome collapse, with large clasts of the
porphyry, passes transitionally into a block-and-ash flow
which in turn passes into laterally into more typical pyroclastic deposits of the Calderones sequence.
Km 0.0: Beginning of the road to El Cubo.
Km 6.6: Road intersection. We take the road to the left,
headed for Peregrina.
Km 9.5: Gate to the Peregrina Mine.
Km 10.1: Road intersection. Turn left, toward Cerro Alto
de Villalpando.
Km 10.4: Road intersection. Turn to the right, going up
a steep hill.
Km 11.4: We will park as far off the road as possible and
walk a short distance up the road for an overview of the
District.
DAY 2: PYROCLASTIC SEQUENCE OF THE CALDERONES
FORMATION AND POSSIBLE SOURCE VENTS.
STOP 2-1: CERRO ALTO DE VILLALPANDO. RING DIKE WITH
CALDERONES VENT FACIES AND A PANORAMIC VIEW OF THE
CALDERONES SEQUENCE FROM ITS PRINCIPAL POINT OF ORIGIN. (UTM 14Q0273434; 2326116)
In this long roadcut just beneath the summit of the hill,
we can observe a large ring dike which is composed of
Calderones tuff and tuff-breccia. The dike cuts a dacitic
to rhyolitic dome that forms the greater part of Cerro Alto
de Villalpando and that we consider part of the Peregrina
dome complex. Overlying the Peregrina dome, and likely cutting it is the Chichíndaro Rhyolite, which forms the
highest part of the hill. The road cuts obliquely across the
ring dike, affording us an excellent exposure of its contacts and its interior along a traverse of several hundred
meters; the true width of the dike is at least 50 meters.
The contact near where we left the vehicles is subvertical
and irregular, with well-developed shear surfaces both
within the dike and in the host rock. In other locations we
have seen breccias and cataclastic rocks in zones up to 8
meters wide along the dike margin. These zones are made
up of well-preserved fragments of the Peregrina dome
rocks in matrices of Calderones tuff. The fragments in the
interior of the dike are heterolithologic, including clasts
derived from the Mesozoic basement, such as phyllite,
argillite, quartzite, meta-andesite and calcareous rocks, as
well as clasts derived from the Cenozoic cover, such as
the GRC and altered rhyolitic to dacitic rocks, presumably from the Peregrina domes. All these clast types
exhibit considerable size variation, ranging from a few
millimeters to several meters in diameter. Extremely
154
large blocks of phyllite of the Esperanza Formation are
present in the dike, suggesting that this formation is present at shallow depths below the surface; one such block
can be seen behind the small white building. This pyroclastic dike probably served ad the principal venting
structure for the Calderones Formation, and it can be
interpreted as a partial-ring fracture bounding a caldera.
Because its outcrop is limited to the northeastern quadrant of the putative circular boundary, a partial or trap
door, morphology is suggested for the Calderones
caldera. After examining the dike, we will avail ourselves of
the excellent panoramic view, looking westward, of the
Calderones sequence. Immediately below the road is an
area with a noticeable reddish brown color, which forms
a band of low, rounded hills in the vicinity of the Tiro de
San Lorenzo. These hillocks correspond to the intracaldera facies of the Calderones sequence, which consists
of a collapse megabreccia made up of enormous clasts of
Esperanza Formation (phyllites and schists) in a scant
matrix of Calderones tuff. In addition to the more abundant Esperanza-dominated megabreccia, there are isolated outcrops of Peregrina-dominated megabreccia. We
will be looking at these outcrops at the next stop. In a
rough way, the map pattern of this megabreccia unit is
parallel to the outcrop of the ring dike, having the same
curvature. In the middle distance, beyond the megabreccia, we
can see a prominent ridge with several summits held up
by thick, apparently massive, layers. The largest and
highest of these summits is Cerro de La Loca, where we
will go for the third stop of the day. The summits are separated from each other by minor faults striking approximately NE-SW, with their northwestern sides downthrown. Cerro de La Loca itself, like the entire La Loca
ridge, is made up of the proximal facies of the Calderones
sequence, with a series of relatively thick and voluminous ignimbrites at the top. The ridge beyond the La
Loca ridge is the La Leona Ridge, whose summit is
capped by a thick section of the Bufa Ignimbrite. The
Bufa is in tectonic contact with the Calderones along the
La Leona normal fault, whose trace follows the base of
the dip slope of the ridge. Unlike most of the normal
faults in the District, it dips to the NE, toward us; the low
hills and ridges on the downthrown block are underlain
by the Calderones and Cedro Formations. We will see
this part of the Calderones sequence (the medial facies),
bounded by the La Leona and El Cubo faults, at the
fourth stop of the day, on a traverse following the Arroyo
de Los Silvestres. Even further in the distance we can see
a series of small tilted mesas on the other side of the La
Leona ridge. These mesas contain the distal facies of the
Calderones, near the village with the same name. The hill
known as Cerro Coronel and the crags called Las Dos
Comadres stand out as landscape features, and we will
visit them at the fifth and sixth stops of the day.
We will return by the same road to the exit from the mine.
Km 12.8: At the first intersection we turn left.
Km 15.2: Mill and flotation plant of the El Cubo Mine.
We take the road to the left, toward the valley where the
Tiro de San Lorenzo is located.
Km 16.3: We will park at the side of the road near a small
ditch.
STOP 2-2: TIRO DE SAN LORENZO: COLLAPSE
(UTM 14Q0273297; 2325698)
MEGABRECCIA
At this stop there are outcrops next to the road and in the
small ditch which show features that we interpret as the
intracaldera collapse megabreccia related to the paroxysmal eruption of the Calderones sequence. The breccia
here includes fragments of various sizes, from at least 10
meters down to a few centimeters; some blocks make up
entire outcrops. By far the most common lithology at all
fragment sizes is phyllite of the Esperanza Formation.
The matrix is difficult to see, because the deposit is
deeply weathered and the fine-grained fraction has mostly been converted to a yellowish brown to reddish brown
soil. Upon careful inspection, however, one can observe
that this soil consists of materials of the same type as the
larger blocks, pulverized to the size of sandy grit or fine
gravel. The deposit as a whole is quite altered by the
action of hydrothermal solutions that permeated it, producing abundant cross-cutting tiny veinlets of quartz and
rendering it susceptible to intense weathering. All of the
larger blocks are pervasively fractured and some fracture
domains show a jigsaw-puzzle pattern. Blocks of the
Esperanza Formation here have a somewhat better developed schistosity than that observed at Stop 1-1. In one
especially large block we see the typical stratigraphy of
the Esperanza Formation: metalava (greenschist), sandstone (quartzite), and metamorphosed shale (schist).
Another large block, internally a phyllite, has a cataclastic margin at its contact with Calderones tuff. Near the
flotation plant that we passed on the road below this stop,
we find another type of component of the megabreccia,
which is a megablock approximately 10 meters long of
finely flow-banded dacitic or rhyolitic Peregrina dome
rock. This block shows a great internal complexity, perhaps inherited from the original dome, which includes a
zone of jigsaw-puzzle fracturing, a highly sheared zone,
and a zone of fault gouge in contact with a clastic (?) or
pyroclastic (?) dike with fragments of many different
types in a matrix of Calderones tuff. Above this dike, in
its apparent hanging wall, we see Peregrina dome breccia
and very fine-grained Peregrina dacite or rhyolite, which
appears to be a devitrified glass.
From this stop we leave the area of the Peregrina Mine
and pass through the same control gate that we crossed
earlier. We take the road back toward Guanajuato, but at
the intersection with the road to the El Cubo Mine we
turn to the left. 0.9 km beyond the turn we will stop next
to Cerro de La Loca in order to climb up it and look at a
section of the Calderones sequence.
STOP 2-3: CERRO DE LA LOCA: PROXIMAL FACIES OF THE
CALDERONES SEQUENCE. TRANSITION FROM THIN IGNIMBRITES
AND INTERBEDDED BRECCIAS UPWARD TO THICK CAPPING IGNIMBRITES. INTERACTION BETWEEN DIKES AND BEDDED
DEPOSITS (UTM 14Q0272690; 2324213)
At the base of the section exposed here is a package of
pale green thin-bedded (5-15 centimeters) deposits composed of fine-grained and well-sorted clastic material
with a considerable ash component. The strata have wavy
shapes, suggestive of gentle folding, draping, or differential compaction over an irregular surface that does not
crop out at the surface, and they lack cross-bedding.
Some beds have rather poorly developed normal graded
bedding, while other beds exhibit flow texture. We think
these beds are best described as a rhythmic sequence of
mixed fall and surge deposits, which were deposited in
very shallow water. The sequence passes upward into a
series of thin to medium-thickness (15–30 centimeters)
pyroclastic flow deposits, which include surges, minor
ignimbrites and breccias of uncertain origin, possibly
medial portions of block-and-ash flows. Here, as in other
parts of the Calderones Formation, almost all the deposits
155
contain fiamme converted to dark green chlorite. There
are many interesting color variations in these beds – in
some layers only the shard-rich matrix and the fiamme
are green; other layers have green fragments in a grayish
matrix, and still other layers have both green lithic fragments and green fiamme/matrix. This series of beds continues for several meters, until a contact is reached with
a package of thin cross-bedded layers accumulated
almost exclusively from pyroclastic surges. The surge
layers are overlain by a series of green thinly-bedded
sandstones which are in turn overlain by another package
of surge-bedded tuffs, rich in large (up to 35 centimeters)
angular lithic fragments of various types (GRC, vein
quartz, granite, etc.).
The pyroclastic rocks that make up the summit of
Cerro de la Loca are a group of thick ignimbrites which
form the culminating sequence of this part of Calderones.
We interpret this entire package as a single cooling unit,
accumulated in at least four pulses of emplacement or
changes in the nature of the pyroclastic density current.
The degree of welding is comparable to the strongest
welding observed in any part of the Calderones; thin sections reveal the presence of spherulites surrounding the
lithic fragments, embedded in a matrix with well defined
eutaxitic texture. Each of the emplacement units has different lithic fragments; for example, the lower most unit
contains sparse small fragments of limestone as well as
many slightly larger angular fragments of Peregrina
dacite, while the upper three units have fragments of
phyllite but lack limestone. The third emplacement unit,
up from the base, has scarce lithics of small size. The
uppermost unit is characterized by an abundance of very
small lithics that varies little from base to top. It is perhaps the unit with the highest lithic content of the four
units that make up this composite ignimbrite. The lithics
are principally of reddish-brown lavas and glassy white
lavas. Pumice is not apparent, but it may have been completely masked by secondary silicification. This uppermost unit has two interesting physical features that are
not present in the lower ones. A prominent pseudo-stratification seems to reflect progressive aggradation resulting from instability of the pyroclastic current (Branney
and Kokelaar, 1992). The second interesting feature is the
presence of broad shallow pits in the uppermost surface,
arranged in a regular grid, which may be the tops of fossil fumaroles. In total, the thicknesses of the four ignimbrite emplacement units add up to about 50 meters.
156
Our interpretation is that these ignimbrites were erupted
during the paroxysmal phase of the eruptions that formed
the Calderones sequence. The lithic-bearing surge
deposits at the base of these summit ignimbrites correspond at least in part to Layer 1 of Sparks and others
(1973); some breccia sheets intercalated with the lower
most of the obvious surge deposits may represent a different mechanism of transport and deposition. A subtle Layer 2a, with shearing features and depletion of larger grains, can be observed at the base of the
first emplacement unit, and Layer 2b makes up the rest of
the unit. Layer 3 deposits appear at the top of the first
emplacement unit as a thin zone of laminated layers,
although it is difficult to confirm its identity because of
the effects of secondary processes, and because it is principally exposed in the middle of the summit cliff. In addition, the putative Layer 3 of the lower unit is directly
overlain by an interval of thin cross-bedded layers, which
almost certainly constitute the Layer 1 surge deposits of
the next emplacement unit. We have not determined the
precise compositions of any of the emplacement units in
this culminating sequence, but given the relatively highsodium content of the plagioclases inferred form petrography, they could well be dacites.
In many places Calderones is cut by andesitic dikes
that are thought to be equivalents, and probably feeders,
of the Cedro Andesite. The interaction between the dikes
and the apparently still water-bearing pyroclastic
deposits of the Calderones locally gave rise to a second
generation of pyroclastic products of phreatomagmatic to
strombolian type. We can observe this phenomenon
about in the middle of the section, where deposits of ash
and breccias lie unconformably over the Calderones
sequence adjacent to an andesitic dike. The dike also produced thermal alteration in the surrounding layers of
Calderones, indurating them and making them more
resistant to erosion than the parts not altered by the dike.
The result of this process is that tabular erosional forms
of the baked Calderones stand up above ground level on
both sides of the deeply eroded dikes. At this locality we will also see a deposit that we
interpret as a lahar. This curious layer covers the stratified layers of Calderones sequence with considerable discordance. The lahar apparently was emplaced long after the
original emplacement of the Calderones (possibly even in
Recent times), and was possibly caused by the conjunc-
tion of intense faulting and abundant rain in this region.
Boulders of the purple dacite of the Peregrina domes
were caught up in remobilized non-indurated fall
deposits or other fine-grained pyroclastic layers, forming
mudflows that flowed along small paleochannels that
may have resulted from ground cracking during small
earthquakes.
After inspecting the Calderones section and other features exposed at Cerro de La Loca, we will take our lunch
break and enjoy the views in all directions, and then will
return to the vehicles. STOP 2-4: ARROYO
DE
LOS SILVESTRES: BOULDER
BEDS,
PYROCLASTIC FLOW LAYERS AND SURGE LAYERS IN THE MEDIAL
FACIES OF
CALDERONES
At this location we are a bit downstream from yesterday’s
last stop (1-7), on an elongate fault block situated
between two large normal faults with opposite dips. As
we face downstream (south), the SW-dipping El Cubo
fault is to our left, and the NE-dipping La Leona fault is
to our right. Thus we are in a major graben which may
have formed before or during the Calderones eruptions,
although both faults have had further displacement after
the Cedro lava flows were emplaced. Evidence for such
movement can be seen in the drag folding of the
Calderones beds along the La Leona fault and in the fact
that the Cedro Andesite was tilted to the NE along the El
Cubo fault. We will walk a short distance down stream in
order to look at a representative portion of the medial
facies of Calderones. There are a number of repetitions of
the typical sequence, which contains (from the base
upward): surge layers of slightly coarser grain size and
more pronounced, larger amplitude duneforms than are
generally seen in Layer 1 deposits. These are overlain by
thick layers (up to three meters) with large boulders in
variable amounts of tuffaceous matrix. These, in turn, are
succeeded by ash-flow deposits with abundant conspicuously flattened and ramped pumices and abundant small
lithics of various rock types; the sequence is capped by a
package of very fine-grained surge layers with unusually
low-angle cross-bedding. As we go downstream (up-section), the boulder layers vary from nearly monolithologic accumulations of Peregrina dome rocks to mixed
assemblages containing more and larger GRC boulders
and a significant quantity of andesite (Mesozoic La Luz
meta-andesite and/or GRC hydrothermally altered
andesite). Near the point where we enter the stream we
can see a number of interesting features associated with
complex branching Cedro Dikes.
STOP 2-5: CALDERONES VILLAGE: DISTAL FACIES OF
CALDERONES FORMATION AND EVIDENCE OF
PHREATOMAGMATIC INTERACTIONS WITH CEDRO DIKES
THE
ITS
In this locality, far from its known sources, the components of the Calderones sequence are more varied than
those we have seen in its proximal and medial facies
(stops 2-1 thru 2-4). The section contains ignimbrites that
are thinner and generally finer-grained (there are some
exceptions) than their counterparts closer to the source
area. There are complicated interactions with the Cedro
dikes; we will inspect the peperites and other interaction
features along the contacts between the rising dikes and
the uppermost, still water-bearing, tuffs of the
Calderones. There are lenses of phreatomagmatic tuffs
interbedded with magmatic tuffs, as well as some
andesite bodies that could be invasive lava flows, near
the dikes.
STOP 2-6: CERRO CORONEL: ENIGMATIC CAPPING UNIT OF
THE CALDERONES FORMATION AND A SUMMARY OF THE MODEL
FOR THE DEVELOPMENT OF THE VOLCANIC PACKAGE WITHIN
THE
DISTRICT
We park the trucks close to the Humboldt shaft and we
will walk upslope looking at the upper part of the
Calderones section. Numerous ignimbrites with spectacularly flattened fiamme in ashy matrix material are
interbedded with increasingly thicker intervals of surge
deposits. This section probably resulted from deposition
in distributive channels and pyroclastic fans in the distal
runout region of the flow currents, and the pyroclastic
material was accumulated in water. The thicknesses of
the layers become smaller near the top of the section. The
final two meters of these deposits beneath the summit
cliff of Cerro Coronel display a wide variety of structures
and a delicate style of the lamination. These layers are a
series of deposits of high-energy surges, that preceded
the emplacement of a large ignimbrite, which caps the
sequence and form pronounced cliffs. These capping ignimbrites form at least two flow units, with notorious
fluid-escape channels with centimeter-scale spacing at
157
their tops. Another characteristic is the high lithic content
of this deposit (more than 30% of the deposit in some
places). As for the lithologic nature of those fragments
(the overwhelming majority are of a moderately to highly welded [?] felsic rock with fine flow banding). The
matrix of the ignimbrite is generally poorly indurated ash
now completely devitried, and very little pumice. We
have an alternative explanation for this capping layer.
With respect to its position along the ‘phreatomagmaticmagmatic spectrum,’ it is possible that it was formed by
phreatomagmatic eruptive processes rather than by purely magmatic processes. Tentative summary of the model for the formation of the
volcanic sequence in the District.
If any one thing is clear, it is that the Calderones is definitely a volcanic unit. It is not a sandstone or a conglomerate made up of detrital fragments derived from older
units, which were transported into the depositional basin
from sources outside the Mining District, solely by
stream action, as was originally proposed when it was
studied by Echegoyén (1970). Precisely how much reworking there was of primary volcanic deposits remains
an important question. Our present model has several elements:
From the first surges of Losero to the final andesitic
flows of Cedro, the volcanic products of the District were
probably associated to a shallow magma chamber.
Magma formed as a consequence of subduction and rose
to a high crustal level in the region during a part of the
early Tertiary. At the time of the volcanism the rate of
regional extension remained high. The eruptions that produced the Calderones were of
several different volcanological types. They include the
growth and collapse of domes, high-mass-flux eruptions
from the ring dike, and minor phreatomagmatic eruptions
related to the rise of the Cedro dikes into the recently
deposited tuffs of the upper member of the Calderones.
Because of the wide variety of eruptive mechanisms, as
well as the variety of environments of transport and
deposition, the deposits are varied with respect to grain
size, fragment shape and angularity, relative proportions
of accidental and juvenile components, textures, structures, thickness of layers, and other parameters.
It seems probable that the Calderones eruptions
began with the rise of the small dacitic domes in the
158
northwestern part of the Peregrina Dome Field.
Successive collapses of these domes produced blockand-ash flows that were emplaced both around the domes
and within recently formed grabens and/or river valleys.
Events in the northeastern part of the Peregrina Dome
Field followed these early collapses without much of an
intervening quiet interval. Initially, some larger and more
silicic domes were emplaced thru and atop rocks that
included the Esperanza Formation, the GRC and the Bufa
Ignimbrite. It is possible that one or more of these domes
contained large roof pendants of the Esperanza
Formation. The “interior” portion of this complex of
domes, wallrocks and roof pendants (interior with respect
to the incipient caldera) collapsed along a fracture coincident with or parallel to the presently exposed ring dike,
producing the megabreccia with its zones rich in
Esperanza fragments or Peregrina fragments, respectively. One good candidate for the scarp produced in this collapse might be the Veta Falla de Villalpando (J.
Echegoyén, pers. comm., 2002). It seems clear that the
ring dike was a very important source of the ignimbrites
of the Calderones sequence, especially those along the
Cerro de La Loca ridge. It is also possible that zones
along the trace of the ring dike formed initially as open
fissures of the extensional type, related to the principal
fracture along which the caldera wall collapsed to form
the megabreccia. Calderones tuffs erupted from other
vents might have filled these zones from above. Almost all the deposition of the Calderones
Formation must have occurred in shallow waters.
Conversion of all the glassy materials to chlorite must
have been caused by the original heat of the pyroclastic
fragments, rather than occurring much later, as a consequence of hydrothermal alteration of cold pyroclasts by
rising hot fluids. It is difficult to know for certain how
many of the individual eruptions of Calderones tuffs
were phreatomagmatic in nature; some of the thin-bedded and very fine-grained deposits may be the result of
phreatoplinian eruptions. Some important things to investigate in the future
include the nature of the southeastern portion of the ring
dike, the nature of the chloritization (isochemical or allochemical) of the glassy fragments, the timing of displacement along the major faults of the District, especially the La Leona Fault, and the details of intrusive structures, compositions and relative ages of individual bodies
within the Peregrina Dome Field. DAY 3: THE FINAL STAGES OF THE SIERRA MADRE
OCCIDENTAL ARC (LATE OLIGOCENE) AND THE BEGINNING
OF THE TRANSMEXICAN VOLCANIC BELT (MIOCENE) IN THE
REGION BETWEEN GUANAJUATO AND SAN MIGUEL ALLENDE
Refer to Figures 8 and 9 for locations of some of the
stops.
We will leave from the lobby of the base hotel.
From here we will hear toward the exit to Juventino
Rosas and San Miguel Allende. The measured distance in
kilometers begins at the Santa Fe de Guanajuato Glorieta
(traffic- circle/rotary/roundabout) in front of the Holyday
Inn Express hotel.
Km 0.0: Glorieta Santa Fe de Guanajuato. We take the
highway to Juventino Rosas and San Miguel Allende.
Km 20.3: We stop at a prominent roadcut. The road here
has narrow shoulders; please be especially careful about
the traffic, which is both fast and heavy.
STOP 3-1: FLOW FOLDS IN THE CHICHÍNDARO RHYOLITE
In this road cut we see intense folding in the midTertiary rhyolites. This deformation is not of tectonic
origin; its origin is syngenetic with the emplacement of
the lava. Because of its high silica content, this lava
must have been very viscous, and upon moving across
the surface of the earth, it must have been deforming in
a complex way. Almost all the original glass is devitrified and hydrothermally altered zones are common.
Tension cracks in fan-like arrangement are sporadically visible in the crests and troughs of some of the flow
folds. Vapor-phase topaz occurs in some outcrops of
the Chichíndaro rhyolite (but not in this particular one)
along some of the flow-bands. These folds and the subvertical flow foliation are very common close to the
centers of emission of the lava, which in this region
commonly forms large endogenous domes and tholoids
or coulees. This particular type of rhyolite is common
and abundant in the extreme southern part of the Mesa
Central. Along with low to medium grade felsic ignimbrites they constitute the principal products of SMO
volcanism in the region (Aranda-Gómez and others,
1983). Labarthe and others (1982) have estimated a
thickness of ~1,000 meters for the mid-Tertiary volcanic sequence (K-Ar: 30–27 Ma) near the city of San
Luis Potosí.
Km 39.0: Intersection of the highway to San Miguel
Allende. We turn north and park the vehicles some five
hundred meters beyond the intersection.
STOP 3-2: PANORAMIC VIEW OF THE VOLCANIC SEQUENCE OF
THE MESA DE SAN JOSÉ DE ALLENDE (LATE OLIGOCENE TO
MIOCENE). HIGH-STAND PORTION OF THE FLUVIOLACUSTRINE
BASIN OF THE RÍO LAJA AND DRAG FOLDING OF BASIN CONGLOMERATES ASSOCIATED WITH NORMAL FAULTING (UTM
14Q0289371; 2311238)
At this locality we will discuss the sequence exposed in
the Mesa de San José de Allende, which can be seen in
the distance to the southeast of us. We will also take a
close look at the highway cut, which exposes sediments
accumulated in the marginal portion of a wide fluviolacustrine basin that extends from here across a broad area
in central Mexico (Figure 3).
The Mesa de San José de Allende is located in the
southernmost part of the Sierra de Guanajuato. In this
part of the sierra, the southern boundary of the Sierra
Madre Occidental province and the Transmexican
Volcanic Belt province overlap, so that it is possible to
find volcanic rocks of both provinces in this sector. Cerca
and others (2000) mapped and dated the units that make
up this mesa in particular and the southern part of the
Sierra de Guanajuato in general. The units exposed in the
mesa section unconformably overlie the Cedro Andesite
(sensu lato), which crops out in small windows at the
edge of the mesa. A K-Ar age of 30.6 Ma was obtained
on andesite collected at this site. Above the Cedro
andesite is a sequence of ignimbrites typical of the SMO;
that is to say, large-volume felsic ignimbrite sheets of
broad lateral extent. Within the mesa section two of these
ignimbrites are exposed, the lower one having a K-Ar age
of 23.1 Ma and the upper one yielding an Ar40/Ar39 age
of 22.4 Ma. The mesa is crowned by andesitic lavas
whose ages (13.2- 13.8 Ma) has been interpreted as the
initial eruptive phases of the Transmexican Volcanic Belt
(Cerca et al., 2000).
In the roadcut you can see a normal fault which puts
the basin gravels into contact with the Cedro andesite.
These gravels are in the hanging-wall block of the fault;
they are relatively coarse-grained and clast-supported.
They show an open synclinal form, which we interpret as
a big drag fold the downthrown block of a down-to-thebasin normal fault. The clasts are derived from the
159
Oligocene volcanic sequence; principally they are rhyolites and andesites ranging from subangular to subrounded in shape. Clasts of andesite predominate close to the
fault trace. Farther away from the fault the most common
lithology, present as large (0.3 meters) to small (0.05
meters) subrounded clasts, is a reddish, nearly aphyric
felsic ignimbrite. Small subangular green chert clasts are
also common. At this site we are in the highest part of the Río Laja
fluviolacustrine basin (Figs. 2, 3), far from the basin axis,
possibly in a depositional environment transitional
between piedmont-type alluvial fans and high-stand
lakeshore zones. We think the gravels were derived by
rapid erosion of the footwall block of the fault and were
deposited very close to their source, at the foot of the
fault scarp.
Km 49.5: Turnoff toward the village of Peña Blanca. We
will park by the roadside.
STOP 3-3: SUCCESSION OF GRAVEL AND SAND DEPOSITS ACCUMULATED IN ALLUVIAL FANS IN THE HIGHER PART OF THE RÍO
LAJA FLUVIOLACUSTRINE BASIN (UTM 14Q0297588;
2314632)
Intercalated gravel and sand deposits are well exposed
along the road cuts of the highway. These deposits were
derived principally from the Oligocene volcanic
sequence. Approximately 50% of the sequence is
formed by lenses of gravel, some of which appear to
have been deposited in well-defined fluvial channels.
Clast size is, on average, about 5 centimeters, considerably less than that observed at Stop 3-2, and clast shape
is more rounded. These possibly represent fluvial
deposits accumulated in braided distributary streams,
which transported distal alluvial-fan deposits down
toward the central part of the basin. The gravels occupied shallow channels, and the sands and silts were
deposited in adjacent pools. Looking in the direction of
Peña Blanca, we can see a resistant unit with crude
columnar joints. This unit is the San Nicolás Ignimbrite
(K-Ar, sanidine, 24.8+/- 0.6 Ma; Nieto-Samaniego and
others, 1996), which is found here intercalated with
gravels similar to those in the road cut.
We continue on the highway to San Miguel Allende. As
we descend toward the Ignacio Allende Dam, we can
160
observe that the average grain size in the gravel horizons
is continually diminishing. At the same time, the relative
amount of sand in the deposit is increasing notably.
STOP 3-4: VIEW OF PALO HUÉRFANO VOLCANO; PRODUCTS
OF PALO HUÉRFANO ABOVE ALLENDE ANDESITE (~11–12 MA)
AND INFERRED AGE OF THE SAN MIGUEL ALLENDE FAULT.
Km 63.0: Village of La Ciénega (UTM 14Q0306846;
2311514). Along this part of the highway there is a small
roadcut where true lake sediments are exposed. West of
the village there are more extensive outcrops of this same
unit. These rocks are made up of fine-grained felsic pyroclastic material and appear as thin-bedded white silts and
clays, in layers 5–10 centimeters thick. Some layers display concentrations of small pumice fragments (lapilli).
Concretions and wavy layers of light brown to yellowish
brown chalcedony are common in this sequence. The
stratigraphic position of these layers in the larger fluviolacustrine sequence of the Río Laja Basin is uncertain.
Some lacustrine sediments are known to underlie and be
interdigitated with products of the Palo Huérfano
Volcano (Stop 3-8) while other lacustrine layers are
known to contain an early Pliocene fossil fauna
(Kowallis and others, 1998). Therefore we conclude that
in the Río Laja sequence, lacustrine deposits must occur
at several levels with different ages. This is what one
would expect, given that the basin was formed as a result
of tectonic activity in several known pulses of different
ages. It is logical to suppose that large lakes periodically
appeared and were destroyed, perhaps in different parts
of a large complex basin.
Km 73.2: Ignacio Allende Dam. The curtain of the dam
was constructed in a narrow canyon excavated in a
sequence of unusually thick andesite lavas. In the road
cuts are outcrops of the Allende Andesite (K-Ar, wholerock, 11.1+/- 0.4 Ma, Pérez-Venzor and others, 1996).
In the canyon of the Río Laja the Allende Andesite is
cut by andesitic domes. According to Cerca and others
(2000), the domes do appear to be in intrusive contact
with the Allende Andesite, so they must be younger than
~11 Ma.
Km 77.2: Intersection of the highways San Miguel
Allende-Comonfort and San Miguel AllendeGuanajuato. Eastward from this point, in the hill with
the microwave towers, the Allende Andesite unconformably underlies products of the Palo Huérfano
Volcano.
Km 81.5: We will park on the right-hand shoulder of the
road for an overview of part of the San Miguel Allende
Volcanic Field.
From this point looking south from the east side of the
highway, one can enjoy a good view of the Palo
Huérfano Volcano and the primary dip of its products
toward the west. Immediately to the right of the highway you can see a mesa tilted to the east. The base of
this mesa contains rocks of the basal Mesozoic complex. The rimrock at the top is the Allende Andesite.
The Allende Andesite was dated by Pérez-Venzor and
others (1996) at about 11 Ma (K-Ar) and by Cerca and
others (2000) at 12.3 ± 0.3 (K-Ar). The andesite is a
microporphyritic rock with a very fine-grained groundmass. The phenocrysts account for less than 5% of the
rock; they are hypersthene and augite, which rarely
reach even a millimeter in length. This contrasts greatly
with the texture and mineralogy of the products of the
Palo Huérfano Volcano, which are always porphyritic
and invariably contain phenocrysts of plagioclase several millimeters long. The precise location of the vent(s)
from which the Allende Andesite was erupted is as yet
unknown.
The interpretation of Pérez-Venzor and others
(1996) is that the Allende Andesite is independent of the
Palo Huérfano Volcano and pre-dates the latest movement on the San Miguel Allende fault. Similar finegrained andesite mesas underlie the La Joya volcano,
which is just to the east of Palo Huérfano. The dip of the
mesa surface suggests that there is a listric component
to the fault motion, which caused eastward rotation of
the andesite along the N-S-striking fault plane. In the
hill with the microwave towers near the village of
Calderón, an unconformable contact between the
Allende Andesite and the overlying Palo Huérfano rocks
is exposed.
Km 89.6: Glorieta/Roundabout at the entrance to San
Miguel Allende. We turn to the east, following the freeway toward Querétaro and Dr. Mora.
Km 93.2: Ignacio Allende Glorieta/Roundabout. We
continue on the highway toward Querétaro.
Km 94.8: Intersection of Highway 111 and the highway
to Dr. Mora. We go straight, toward Querétaro.
Km 110.0: We park here for an excellent overview of the
central part of the San Miguel Allende Volcanic Field.
STOP 3-5: JUNCTION OF THE ROAD TO GUADALUPE TAMBULA:
PANORAMIC VIEW OF THE VOLCANOES PALO HUÉRFANO, LA
JOYA, AND THE CERRO COLORADO DOME, AND ALCOCER-LA
ESTANCIA FAULT SYSTEM. From this point on the edge of the highway, looking
toward the southeast, we have a good panoramic view of
the northeastern flank of the Palo Huérfano Volcano and
the Cerro Colorado Dome (Figure 9). The most notable
feature is that part of the lava flows from Palo Huérfano
dip to the south, back toward their source vent area.
Pérez-Venzor and others (1996) show on their map a system of normal faults with average N80E strike and blocks
successively downthrown toward the north (Figure 8).
These authors think that movement along listric fault
planes, which forced the layers to rotate counterclockwise (as seen from the east), caused the anomalous dip of
the lava flows. The maximum age of the N80E fault system is given by the age of the volcano (~11 Ma). These
faults are parallel to the system of faults known as the
Chapala-Tula system, which crosses the TMVB south of
this area (Figure 2).
The Cerro Colorado dome is a volcano that predates
Palo Huérfano. It is composed of dacites and has a radiometric age (K-Ar, plagioclase) of 16.1 ± 1.7 Ma (PérezVenzor et al., 1996). The age relationship is apparent on
aerial photographs and on the geologic map, because
Cerro Colorado acted as a topographic barrier, which
deflected the lava flows emitted from the crater of Palo
Huérfano.
Looking toward the SE we see the neighboring volcano, La Joya, which is very similar in age and morphology to Palo Huérfano. Valdéz-Moreno and others (1998)
provide a good report of the geologic evolution of the
volcano, as well as two K-Ar ages (Figure 10). The history of La Joya is similar to that of Palo Huérfano and
begins with an andesitic dome called El Maguey.
Although this dome was not dated radiometrically, it
seems likely that it was contemporaneous with the Cerro
Colorado dome that underlies Palo Huérfano. In fact, the
domes are very close to each other. Later fine-grained
andesitic lavas were emplaced in the area northeast of La
Joya; these lavas resemble the Allende Andesite petrographically and may well be equivalent to it. ValdézMoreno and others (1998) called these lavas “Older
Andesite.” The products erupted by La Joya were
deposited on top of these older andesites. The initial
161
flows of La Joya package are andesitic porphyries whose
most outstanding characteristic is the presence of light
green enclaves which are interpreted as glomerophenocryst clusters that formed part of the ‘crystal mush’
attached to the walls and/or roof of the subvolcanic
magma chamber. These lavas yield an age of approximately 10 Ma. Above these enclave-bearing flows are
more andesite lavas, which built up much of the lower
part of the volcano. The lavas that form the principal
body of the La Joya Volcano are dacites with K-Ar ages
around 9.9 Ma. Finally, the southern flank of the volcano
was partially covered by mafic andesites of 6.2 Ma, produced by a field of cinder cones located to the south of
the volcano. Interesting features of La Joya and Palo
Huérfano are the broad depressions in their summit
regions. These physiographic forms are the result of erosion of the volcanic craters to a broad circular to elliptical depression roughly 3 kilometers wide and 5 kilometers long. This deep erosion was favored by intense alteration resulting from fumarolic activity (possibly continuing for a long time after the final stages of construction of
the volcano). Further along the route we will have opportunities to observe the dacitic flows, and we will go up
into the eroded ‘cirque-like’ summit area of La Joya.
Km 121.0: Village of La Monja. North of the highway
there is a small ledge of material from which andesite has
been quarried. We will park here for our next stop.
STOP 3-6: VILLAGE OF LA MONJA: BASE OF LA JOYA
VOLCANO AND OUTCROP OF THE LOWER ANDESITE WITH GREEN
ENCLAVES
Valdéz–Moreno and others (1998) considered these rocks
as lava flows entirely pre-dating La Joya, but now we
believe that in reality they are the first flows to come out
from the volcano. From this locality Valdéz-Moreno and
others (1998) obtained an Ar40-Ar39 age of 10.4 to 10.9.
The rock is dark gray to black on a fresh surface, with
porphyritic texture and aphanitic groundmass. The phenocrysts form 15-20% by volume of the rock; they are
plagioclase and orthopyroxene. One of the most noticeable features of the outcrop is the presence of fractures
that divide the lava flow into rough and irregular
“sheets.” In other locations, such as the village of Santa
Inés, the spacing between the fractures is so regular that
it permits quarrying of this rock type as a construction
162
material which is used as an ornamental facing stone on
buildings and as a fancy paving stone. Seen at close
range, the andesite contains yellow and green enclaves,
which resemble medium-grained sandstones. Their sizes
vary from 5 to 30 centimeters in diameter, and their form
is noticeably rounded. We think these inclusions were
once part of the crystal mush attached to the walls/roof of
the magma chamber; such features are common in viscous lavas such as those that form domes. The enclaves
are composed of crystals of pyroxene and plagioclase in
an interlocking texture, with intersertal glass (ValdézMoreno and others, 1998). Lavas at a higher level on the
volcano (the Tambula Dacite of Valdéz-Moreno and others, 1998) contain similar enclaves.
At La Monja we will take the road toward La Barreta and
La Joya in order to go up to the summit of the volcano.
We will be ascending through several dacite flows with
autobrecciated bases. The dacite is light grey, with phenocrysts of plagioclase and mafic minerals. The sequence
begins with pyroxene andesite at the base (e.g., La Monja
Andesite), and changes upward to hornblende dacite at
the summit (i.e., the Pinalillo Dacite). The lavas of this
volcano certainly were very viscous; they were emplaced
as lobate flows with very high aspect ratios. Apparently
the viscosity was increasing with time, culminating with
short flows of the coulee type.
Km 129.4 Summit of La Joya volcano.
STOP 3-7: EROSIONAL ‘CALDERA’ AT THE SUMMIT OF LA JOYA
VOLCANO
Once we reach the summit we can see the broad cirquelike erosional form, which was developed from the original eruption crater. The original rock has been totally
argillized to a slippery, fragile, easily eroded yellowish
clay. In a few places vestiges of the original dacite can
still be seen, although it is highly altered by hydrothermal
activity close to the conduit of the volcano. Bit by bit the
original crater was widened and degraded by the erosion
of this argillized material until it ended up as the broad
depression we now see.
From this point we begin the trip back to the city of
Guanajuato. In San Miguel Allende we will make the last
two stops of the day.
Km 165.6: Ignacio Allende glorieta/Roundabout. The
place where Highway 111 joins the freeway south of San
Miguel Allende. On one side of the mall where the
Gigante supermarket is located, a paved road takes off.
We turn to the right and follow the signs leading to the
park called El Charco del Ingenio. Km 167.6: Entry gate of the botanical garden El Charco
del Ingenio. We will leave the vehicles and follow the
pathway that leads to the canyon located downstream
from the El Obraje Dam.
STOP 3-8: CHARCO DEL INGENIO: EL OBRAJE IGNIMBRITE
(~29 MA) AND LAHAR PRODUCED BY PALO HUÉRFANO
VOLCANO RESTING ON ASH AND/OR EPICLASTIC/VOLCANIC
MATERIAL DEPOSITED IN A LAKE (UTM 14Q0320281;
2314024)
We follow the trail toward the wall of the dam and the
canyon. At the mouth of the dam, and along the north
wall of the canyon the rhyolitic El Obraje Ignimbrite
crops out (Pérez-Venzor et al., 1996). This is an extensive
lithologic unit, exposed principally to the north of Palo
Huérfano, in the footwall block of the San Miguel
Allende fault. At the base of the volcano itself, only two
small outcrops of this ignimbrite have been recognized
(Figure 9), resting unconformably on the basal Mesozoic
complex.
In this locality we estimate that the ignimbrite has a
minimum thickness of ~100 meters. It shows a rough
zonation, evidenced by a change in the fracture density,
and the degree of welding, as well as the presence of several zone s of flattened lithophysae partially filled by
quartz and chalcedony. The texture of the rock is porphyritic, with 25–30% of phenocrysts (quartz, sanidine,
sodic plagioclase, and totally altered (?)hypersthene) set
in an aphanitic matrix. We think that this unit is a good
candidate for an ignimbrite that conforms to the progressive aggradation model of emplacement; i.e., it is composed of several packages which are separated by shear
surfaces, presumably reflecting periodic changes in the
flow regime during an essentially continuous eruption.
Discrete shear surfaces may reflect sudden, temporary
increases in boundary-layer shear between particulate
and non-particulate components of the flow (Branney
and Kokelaar, 1992, and Branney, in press). Although
this unit has neither abundant lithics nor obvious
pumices, it does have some layers characterized by sub-
tle development of ramp structures whose orientations
suggest a transport direction from north-northwest to
south-southeast. Pervasive devitrification of original glass and local
silicification partly obscure original textures and structures within this unit.
The El Obraje Ignimbrite is very similar to volcanic
rocks with radiometric (K-Ar) ages of ~32-27 Ma, which
form the majority of the outcrops in the region located
between San Miguel Allende and San Luis Potosí.
Along the path situated on the south wall of the
canyon are some outcrops of an epiclastic/volcanic
deposit which rests on crudely stratified white pumicerich siltstones derived from a tephra fall or some other
volcanic source. On the basis of their appearance and
composition (andesitic), these aquagene pyroclastic
deposits are thought to come from Palo Huérfano
Volcano (called the San Miguel Allende “tuff” by PérezVenzor and others, 1996). The fine-grained sediments
here are very similar to clastic material of lacustrine origin (see description of the sedimentary sequence exposed
in roadcuts near the village of La Ciénega), although we
have no certainty that they are correlatable, nor even that
in this locality they were deposited in a lake of any depth. Overlying this deposit is a thick (several tens of
meters) lahar with large andesitic clasts in a muddy
matrix, probably produced by Palo Huérfano Volcano.
Exposures of the base of this lahar and the top of the
white silt provide numerous examples of soft-sediment
deformation in the lower unit; asymmetric load casts and
shear features suggest that the lahar was emplaced from
south to north.
The marked contrast in the thickness of the El
Obraje Ignimbrite in the regions located to the north and
to the south of the N80W system of faults can be interpreted as evidence that this system could have been active
before the formation of the Palo Huérfano Volcano. The
activity is constrained to be in the interval between the
emplacement of the ignimbrite (~29 Ma) and the age of
the volcano (≤11Ma). This presupposes that the ignimbrite also was deposited in the zone presently occupied
by the volcano, but that after the faulting took place it was
almost totally eroded from the upthrown (southern)
block. An alternative interpretation is that at the time of
emplacement of the ignimbrite, the zone now covered by
the volcano constituted a basement high, which could not
be surmounted and covered by the pyroclastic flow.
163
We return to the vehicles to continue the excursion.
Km 169.6: Ignacio Allende glorieta/roundabout. We take
the freeway south from San Miguel Allende. At the roadcuts in the high stretch of this highway you can see volcaniclastic deposits of andesitic composition. These
came from Palo Huérfano and overlie rocks of the basal
complex and isolated remnants of the El Obraje
Ignimbrite.
Km 171.6: We will pull off the road and park in a small
turnout, which affords an excellent view, to the north, of
the scarp of the San Miguel Allende fault.
STOP 3-9: HIGHWAY
ROADCUT ON THE FREEWAY: MESOZOIC
MARLS AND CALCAREOUS SHALES, MINOR THRUSTS, AND THE
AFOREMENTIONED PANORAMIC VIEW OF THE SCARP OF THE
SAN MIGUEL ALLENDE
2311984)
FAULT
(UTM 14Q033318796;
At this locality we have several cuts that show in a spectacular way some intensely deformed marine sediments
of the basal Mesozoic complex. The sequence exposed
here consists of marls and argillaceous limestones,
weathered to brownish yellow and reddish brown hues.
In some places we can see the color of the fresh rock,
which varies from medium grey to black, possibly
becoming darker as the presence of organic material
becomes more abundant. Also present here are numerous
small veinlets of calcite and gypsum. The sequence is
well-stratified, but in many places the lateral continuity
of layers is interrupted. Also, the rocks display a dense
fracture pattern of tectonic origin (fissility, or spaced
cleavage), that cuts across the layering at a relatively low
angle. Looking closely at some parts of this cut, you can
see overturned (many fully recumbent) folds on the order
of a few meters, with subhorizontal axial planes. In general, the cleavage is axial-planar to these nearly isoclinal
similar folds, and there are numerous examples of cleavage refraction between layers. In some layers original
sedimentary (fluvial?) cross-bedding is preserved, being
only slightly deformed during the development of the
fracture cleavage. In addition to the meter-scale folds,
there are zones of microfolds (decimeter-scale) and
crenulations (intense folding at the centimeter scale or
less) in the clay-rich layers. Gypsum veinlets commonly
are parallel to the fracture cleavage, and they show some
tendency to be especially close-spaced in the hinge zones
164
of the larger folds; calcite veins typically form perpendicular to the axial-planar cleavage, along extension fractures, and in many places can be seen to cut the veinlets
of gypsum.
Viewed at a distance, the cut shows the traces of
several faults with relatively low dips. This is evident
from the fact that the layering is cut and displaced.
Following with one’s gaze along these fault traces, it is
possible to observe that they do not cut all the layers
exposed in this outcrop, but are lost from view by passing
into bedding-plane-parallel surfaces, or, alternatively,
they may be buried by overlying layers which are apparently not offset. It is not possible to make any visual correlation of the displaced layers so as to establish whether
the movement on the faults is actually normal or reverse.
The interpretation that this is a zone of thrusting with
apparent offset toward the northeast is in agreement with
the type of compressive deformation that the rocks in
these outcrops show as a whole. We think that the fault
traces are “lost” in this cut because the motion was
accommodated in part along the bedding planes. The very
nature of these rocks, extremely clay-rich, and the presence of small quantities of gypsum in the sequence,
would facilitate this type of mechanical response to a
compressive stress at high crustal levels. Viewed in a
more regional way, this outcrop appears to be the footwall
block of a major thrust, since rocks of the Guanajuato Arc
crop out in the arroyo located immediately to the south of
this cut. They are possibly covered by later marine sedimentary rocks; Chiodi and others (1988) recovered an
Aptian-Albian ammonite from this locality.
From this overlook we have a view of the historic
center of the city of San Miguel Allende, which is constructed on the fault scarp of the same name. This is a
notable topographic feature, which can easily be seen in
the region north of the Palo Huérfano Volcano. From the
point where we made Stop 3-5 (the intersection with the
road to Guadalupe de Tambula) Highway 111 is constructed on a small mesa which terminates abruptly upon
reaching San Miguel. The base of the Ignacio Allende
Dam is located on the downthrown block. The San
Miguel Allende fault has an approximately N-S strike,
the sense of motion is normal, and its trace is buried by
the products of Palo Huérfano. Thus, the age of its latest
movement has to be earlier than 11 Ma.
End of the road log for Day 3. We return to Guanajuato.
DAY 4: GEOLOGY OF THE BASAL MESOZOIC COMPLEX
We will leave from the lobby of the Hotel Parador San
Javier, travelling north on the highway toward Dolores
Hidalgo. Refer to Figure 5 for locations of stops.
Km 3.5: Road to the face of the Esperanza Dam and the
Los Insurgentes campground. We turn left. About 250
meters in from the highway is the dam, which was constructed around 1894 by Ponciano Aguilar, a well-known
engineer from Guanajuato (aguilarite, a silver selenoid
that was first described here in the Guanajuato District, is
named in his honor). This reservoir is used to supply
water to a good part of the cit. At the south end of the dam
can be seen submarine lavas of andesitic composition,
which are intercalated with the limestone and pelitic sediment sequence of the Esperanza Formation.
Km 4.6: At the entrance to the campground we will leave
the vehicles and will walk along the road for about 300
meters.
STOP 4-1: CONTACT
ALONG THE FAULT
(VETA MADRE)
BETWEEN ROCKS OF THE VOLCANOSEDIMENTARY AND VOL-
SIERRA DE GUANAJUATO.
THE LA PALMA DIORITE: AN EXAMPLE OF THE INTERNAL COMPLEXITY OF THE BASEMENT UNITS (UTM 14Q0265112;
2329247 TO 14Q0264645; 2329081)
CANOPLUTONIC SEQUENCES OF THE
The road cuts at the beginning of this traverse expose
rocks of the volcanosedimentary sequence of the Arperos
Basin. These rocks, which belong to the member less rich
in limestone of the Esperanza Formation of Echegoyén
(1970), has a more chaotic style of deformation than that
observed at Stop 1-1, with disharmonic folding and
chevron-style folds. This outcrop has veinlets of quartz
and/or calcite filling extension fractures, and well-developed shear zones which separate individual packages of
folds. We attribute these fractures and zones of extension
to the proximity of this outcrop to the great normal fault
now occupied by the Veta Madre, which in this part of its
trace is a very broad zone of offset and mineralization. A
band of cataclasites or fault breccias, strongly oxidized,
separates the volcano-sedimentary sequence from the La
Palma Diorite, which is one of the units of the volcanoplutonic sequence (Figure 11).
La Palma Diorite is a very complex lithologic unit,
which at this locality is dominated by massive microdi-
orites and a group of dikes, some of them felsic tonalites
and others diabases. In other outcrops of the La Palma
Diorite we see magmatic breccias and zones of pegmatitic diorite. The great abundance of dikes with chilled
margins and the variable textures of the host rocks are
distinctive features of this unit.
We will return to the village of La Valenciana and from
there we will take the road to Cerro El Cubilete.
Km 0.0: Church at La Valenciana. The road to El
Cubilete begins here.
Km 0.25: On the left is the road to the La Valenciana
Mine, which is one of the most famous bonanzas of the
District. A few hundred meters from the road junction it
is possible to observe a majestic old building, which is
the ruin of the great benefaction plants constructed by the
Spaniards. The road here crosses the La Palma Diorite.
From this point on, until we arrive at Cerro El Cubilete,
the route provides many cuts with excellent exposures of
the rocks of the Guanajuato Arc. Km 0.6: Guadalupe Mine. This building, with its buttresses in the form of stylized elephants, is in the process
of reconstruction, with a high priority being placed on
preserving as much as possible of the historic materials
and construction.
Km 3.0: On the left is the village of Llanos de Santana,
and some 200 meters to the right is the shaft of the San
Elias Mine. The shaft is in the hanging wall block of the
Veta Madre Fault.
Km 4.7: On the right is the road that leads to the La
Cebada Mine, a property of the Peñoles Group. It is the
farthest west of the active mines of the Veta Madre system. From this point, the road follows approximately
along the contact between the dike complex emplaced in
the La Palma Diorite and the Cerro Pelón Tonalite.
Km 6.6: On the right is the road which leads to the village of Mesa Cuata, and which passes through the summit of Cerro Pelón, the type locality of the unit of the
same name. In this locality there is a plagioclase granite
cut by dikes. Along the road one can see alternatively outcrops of the granite and outcrops of the La Palma Diorite.
In some places hydrothermal alteration can be noted, as
well as intense weathering. Despite these features, evidence of deformation of these rocks can be seen as well.
Km 11.0: The road crosses the Cerro Pelón Tonalite. On
the left, in the distance, is El Cubilete. 165
Km 12.3: To the west is the old mining town of La Luz.
At this site was the first discovery of gold and silver in
the Sierra de Guanajuato, which dates from the sixteenth century. The road continues on the Cerro Pelón
Tonalite.
Km 12.6: We will park on the shoulder to inspect this
outcrop.
STOP 4-2: CERRO PELÓN TONALITE CRISS-CROSSED BY DIABASE DIKES; FAULTING AND ASSOCIATED DRAG FOLDS (UTM
14Q0261867; 2331021)
In this roadcut the exposure shows the tonalite and diabase dikes which make a spectacular geometric pattern.
These dikes are similar in lithology to the diabases
observed cutting the La Palma Diorite at Stop 4-1. They
are interpreted as feeder dikes for the abundant pillow
lavas of the volcanoplutonic sequence. However, the
nearly horizontal position of these mafic bodies seems
more consistent with a group of deformed sills. At this
locality both the sills and their host rocks are cut by several small faults, and one can try to use the geometry of
the drag folds and other features to decipher the sense of
motion on the faults.
Km 13.9: We will pull off the road for a brief inspection
of the outcrop.
STOP 4.3: CERRO PELÓN TONALITE CUT BY DIABASE DIKES IN
UNDEFORMED CONJUGATE SETS (UTM 14Q0260277;
2331960)
In this road cut we again see the Cerro Pelón Tonalite, cut
by a good number of diabase sills, but the structural style
of these is quite different. In contrast to the sills at Stop
4-2, where we see drag folds and irregular margins, this
group has planar margins, without evidence of folding.
They commonly form conjugate pairs around sigma-1.
These two stops, with their contrasting structural styles,
suggest that the deformation, which affected the Cerro
Pelón Tonalite after the emplacement of the diabasic
dikes and sills was heterogeneous on a small scale, being
brittle in some places and brittle-ductile in others.
Km 14.2: On the right is the church of the village of La
Luz, and in front of us is the astronomical observatory of
the University of Guanajuato.
166
Km 14.8: At the bottom of the gorge, on the right, is the
Bolañitos Mine, a property of the Peñoles Group. It is
one of the few active mines on the La Luz vein system.
Km 15.5: On the left is the abandoned Golondrinas
Mine, also a property of Peñoles. Some 200 meters south
of the turnoff is the observatory. At this site we have the
contact between the La Palma Diorite, with its dike complex, and the pillow lavas of the volcanoplutonic
sequence. From this point on, the road crosses both pillowed and massive lavas.
Km 16.1: On the right is the road to the Bolañitos Mine;
in the distance, in the same direction is Cerro El Gigante,
which is covered by mid-Tertiary volcanic rocks.
Km 17.2: The road we are following joins the old road to
the village of La Luz.
Km 17.7: On the right is the road to the Asunción Mine
(Peñoles); close to us is a small ridge on top of which are
some antennas. The hill is composed of rocks considered
to belong to the La Palma Diorite (the dike complex is
prominent here). This outcrop is interpreted as a klippe.
The dike complex rests on the pillow lavas of the volcanoplutonic sequence, and it is unlikely that the section
is overturned. The klippe is cut by the master fault of the
La Luz vein system, which crosses the road at this point.
Km 18.8: On the left is the cemetery of La Luz, which is
located on another small klippe. In front of us is Cerro El
Cubilete, and on the right is the El Bajío Plain of
Guanajuato. The road continues on volcanic rocks of the
La Luz Formation (basalts and andesites), which are
almost completely free of vegetation.
Km 20.3: We will stop to observe the low outcrops on
the left. STOP 4-3: SUBMARINE
LAVAS OF THE VOLCANOPLUTONIC
SEQUENCE(UTM 14Q0256236; 2328009).
The objective of this stop is to show the submarine lavas
with pillowed basalts of the La Luz Formation
(Echegoyén, 1970), their normal stratigraphic position
(evidenced by the geometry of the pillows) and the variable nature of the deformation which affects them. In
general, deformation is controlled partly by the lithology
and partly by discrete shear zones. A this stop, the penetrative deformation is well developed in the originally
hyaloclastitic matrix between the pillows of lava.
Walking along the road toward the village of La Luz, one
can see that the basalts were also affected by deforma-
tion, which converted them into chlorite schists. In some
places the basalt is better preserved within a matrix of
mylonitic schist. In the less deformed pillows one can see
vesicles around their margins. Although their lithology is
in broad aspect similar to that of the lavas and tuffs intercalated in the Esperanza Formation (Stop 1-1.) in the volcanosedimentary complex, these lavas of the La Luz
Formation are thrust over the volcanosedimentary rocks.
We think their environment of accumulation was closer
to the arc axis, and therefore constitute a separate formation. Km 20.5: On the right there is a dirt road heading toward
the village of Los Lorenzos. We can still see Cerro El
Cubilete in front of us. From this point on, the La
Valenciana-El Cubilete road passes through one of the
most evolved facies (in terms of magma composition) of
the whole Mesozoic volcanic sequence. This sequence
includes dacites and rhyodacites (including some keratophyres), as well as pyroclastic rocks of the same composition.
Km 23.3: We will pull off the road here to take a look at
the long outcrop on the right.
STOP 4-4: PYROCLASTIC
ROCKS METAMORPHOSED TO CHLO-
RITE
INTENSELY
SCHISTS
AND
DEFORMED
(UTM
14Q0254905;2325895).
Along several tens of meters the road here crosses chlorite schists which are interpreted as metamorphosed
pyroclastic rocks (probably andesitic to basaltic tuffs
before metamorphism). They were subjected to multiple
episodes of deformation. The early schistosity was folded by a later compressional event. Most of the microfolds
in this area have axial planes that strike NW-SE. If this
section has not been overturned, the schistosity indicates
right-lateral shear, or north-over-south. The most spectacular and elegant feature in this outcrop is the marked
contrast in deformational style between two types of
tuffs. The mafic ash-rich tuff developed a close-spaced
cleavage, crinkle-folds and small anastomosing veinlets
of quartz in the cores and hinge zones of the microfolds.
Its neighbor, on the other side of a small thrust, is a tuff
(probably andesitic) with relatively abundant crystals. It
responded differently to the deformation, with a parallel
style of folding rather than the similar-fold style of the
fine-ash tuff.
Km 24.8: Road crossing. On the left is the road that goes
down to the highway at Silao; to the right is the road that
climbs to the summit of Cerro El Cubilete and the sanctuary of Cristo Rey.
Km 24.8: The road crosses unconsolidated gravels and
sands of Tertiary age, later than the Late Oligocene and
earlier than the mid-Miocene, and goes up toward the
monument of Cristo Rey through andesites of midMiocene age (the El Cubilete Andesite, K-Ar, wholerock, ~13.5 Ma: Aguirre-Díaz and others, 1997).
Unfortunately a stone wall constructed along the edge
of the road prevents our seeing the gravels. We will
visit an outcrop (Stop 4-6) of these gravels on the road
which goes to Silao. At the very top of the hill a chapel
and the Cristo Rey monument were constructed. This
sanctuary is considered an offering in honor of those
who died during the Cristera Revolution, which
occurred en the decade between 1920 and 1930 in this
region.
STOP 4-5: CERRO EL CUBILETE: A PANORAMA OF THE SIERRA
GUANAJUATO AND ITS SURROUNDINGS (UTM
DE
14Q0253836; 2325061).
Views around the Sanctuary of Cristo Rey. From this
point, at 2590 meters above sea level and 700 meters
above the El Bajío Plain, one can see the principal features of the Sierra and its surrounding areas:
To the WNW is El Bajío, with the city of León at its western end. More to the right is the NW part of the Sierra de
Guanajuato; there the vocanoplutonic sequence of the
Guanajuato Arc and the volcanosedimentary sequence of
the Arperos Basin have been intruded by an extensive
early Tertiary (53+/- 3 and 51+/- 1 Ma, K-Ar, biotite:
Zimmerman and others, 1990) batholith known as the
Comanja Granite. Close to us are outcrops of a massive
diorite exposed near the village of Tuna Mansa. These
plutonic rocks were thrust over the submarine lavas
which we saw at Stop 4-3, near Cerro El Cubilete. This
tectonic contact is visible in the downthrown block of the
Villa de Reyes graben.
To the north are the two hills El Gigante and La Giganta,
which are crowned by andesites of Miocene age, similar(?) to those here at El Cubilete, but of different ages.
The depression to the west of El Gigante and La Giganta
is the Villa de Reyes graben, a mid-Tertiary structure
with a N45E orientation. The graben has about 150 kilo-
167
meters of length and in the Sierra de Guanajuato terminates against the El Bajío Fault (Figure 3). The master
fault on the east side of the graben puts into contact the
massive diorite and the submarine lavas. At the center of
the graben is the village of Arperos, and farther north,
inside the graben, is the Sierra El Ocote, a rhyolite dome
with tin and topaz (Figure 2). The Villa de Reyes graben
is the northwestern tectonic limit of the Veta Madre.
To the east of Cubilete is the city of Guanajuato, which
was constructed in a depression bounded by the three
major faults of La Aldana, Veta Madre and El Bajío. The
high area to the northeast of the city of Guanajuato is the
Sierra de Santa Rosa, covered principally with midTertiary felsic volcanic rocks. Beyond the Sierra is the
Palo Huérfano Volcano, which we saw yesterday. San
Miguel Allende sits north of the volcano. In the same
direction is a depression oriented ENE-WSW, known as
the La Sauceda graben, a late Cenozoic structure which
forms the southeastern boundary of the Veta Madre system and also the southeastern boundary of the Sierra de
Guanajuato.
To the southeast on a clear day one can see the opposite
end of the El Bajío Plain and some of the volcanoes of the
Transmexican Volcanic Belt, such as La Gavia, Culiacán
and La Batea. To the south bordering the El Bajío depression is the fault
zone of the same name. The trace of the fault is close to
the base of the mountains, near the village of La Ermita.
This fault was active in the late Cenozoic and caused relative sinking of the Bajío block relative to the Sierra de
Guanajuato. This is demonstrated by the sequence of
gravels and the andesite of El Cubilete, which has been
displaced ~600 meters upward with respect to its counterpart on the downthrown El Bajío block. An estimate of
the rate of displacement on the fault, assuming that the
600 meters were accumulated in the last 13.5 million
years is 0.04 millimeters per year. STOP 4-6: CERRO EL CUBILETE: GRAVELS AND ANDESITIC
LAVAS OF TERTIARY AGE CROWNING THE MESOZOIC BASEMENT
OF THE SIERRA DE GUANAJUATO
At the base of the Cristo Rey monument are exposed
unconsolidated fluviolacustrine deposits composed principally of clasts of volcanic rocks derived from the midTertiary units of the Sierra de Guanajuato, which include
ignimbrites, rhyolitic dome and flow rocks, and
168
andesites. Fragments derived from the basal Mesozoic
sequence are relatively rare. All the clasts in this deposit
are well rounded, and they are supported by a matrix of
coarse sand, fine gravel and silt. The deposit has a poorly developed stratification, marked by changes in grain
size in the finer-grained clastic deposits, and by lenses of
coarse gravel. These deposits reach an aggregate thickness of 150 meters in their thickest part, but they thin
against the Mesozoic rocks, indicating that this is the filling of a paleochannel formed by erosion of the Mesozoic
rocks. No fossils have yet been reported in this sedimentary sequence, but it is inferred to be early to middle
Miocene, based on the fact that the youngest lithic clasts
that it contains are from late Oligocene (around 27–30
Ma), and that this deposit is covered by andesite lavas of
mid-Miocene age.
Resting unconformably on the fluviolacustrine
deposits is a thick subaerial andesitic lava flow, which
caused hydrothermal alteration at the contact with the
gravels and sands. The base of the flow is an autobreccia
which changes upward to a zone with well-developed
columnar joints. Above this zone, without a marked interruption other than a prominent horizontal joint, is another
thick andesite, very similar in texture to the earlier one.
This one has a base with intense subhorizontal platy jointing which passes upward into a thick columnar zone in the
middle part, and culminates with a zone of subhorizontal
platy jointing in the uppermost part. In total the andesite
sums up to 70 meters of thickness and we consider that it
could be formed by two flows. An alternative interpretation is that the El Cubilete Andesite is actually one very
thick (intracanyon?) flow with complex colonnade-andentablature structure. Apparently the lava flowed down a
paleochannel similar to that in which the underlying gravel and sand were deposited, but we do not know from this
small remnant whether or not the lava flow itself was confined by canyon walls. We think the El Cubilete Andesite,
like the volcanoes of the San Miguel Allende Volcanic
Field (Day 3), is associated with the earliest phases of volcanism of the Transmexican Volcanic Belt.
From this stop we will head for the Bajío Airport, where
some members of the group will board their flights to
return home. Those participants who can remain in
Guanajuato will be able to visit outcrops of the Comanja
Granite with us.
STOP 4-7: THE COMANJA GRANITE: A PALEOCENE BATHOLITH
WITH ABUNDANT TOURMALINE IN MAGMATIC, PNEUMATOLYTIC
AND HYDROTHERMAL PARAGENESES (UTM 14Q0246333;
2338599)
In the vicinity of El Rancho Los Alamos, we will visit
an outcrop of the Comanja Granite. This part of the
batholithic body is characterized by many large phenocrysts of alkali feldspar with prominent Carlsbad
twinning; the coarse-grained groundmass has abundant
quartz, a bit of plagioclase (probably albite), and a trace
of biotite. There are some parts of this outcrop that have
rough alignment of the phenocrysts, as if they were flow
domains alternating with static domains, or possibly this
is a secondary structure imposed upon the granite. There
is a rhombic pattern of fractures, some of which are
filled by aplite dikes. There are also enclaves a few centimeters in diameter of very fine-grained material within the coarser-grained granite. Tourmaline is present in
miarolitic cavities, in which it has the common doublyterminated form; in the pegmatitic domains within the
granite it has other habits. This mineral is also present as
a magmatic phase, in acicular form, and in hydrothermal
breccias in massive microcrystalline to cryptocrystalline
form. One can see, about halfway up the slope of this
hill, a wide brittle shear zone, which has various mixtures of massive tourmaline with granite fragments of
different sizes and degrees of comminution. These mixtures range from jigsaw-puzzle arrangement of larger
granite fragments and relatively little tourmaline, to a
distinctive “rosette” pattern of rounded granite clasts in
a matrix of tourmaline and very finely comminuted
granitic material and secondary silica, to massive tourmaline with little or no granite. From time to time traces
of pyrite and other sulfides can be seen in these shear
bands. It seems probable that a careful search of this
outcrop and its surroundings might reveal some small
roof pendants of rocks of the volcanosedimentary complex. UNIVERSIDAD NACIONAL AUTÓNOMA DE MÉXICO, INSTITUTO DE GEOLOGÍA, PUBLICACIÓN ESPECIAL 1 GEOLOGIC TRANSECTS ACROSS CORDILLERAN MEXICO

The geology around the city of Guanajuato is specially interesting

Transcription

Similar documents

San Miguel de Allende - Carnegie Museums member

INTRODUCTION Few countries are blessed with as much diversity

Rock Star 101

Cuetzalan - Ketzaltour

Fire History of Conifer Forests of Cerro El PotosÍ, Nuevo LeÓn, Mexico

Activities and challenges to face the detection of volcanic ash

Volcano Trail: Introduction

Risk management COBRA Network Environmental liability,

z Frank Berbuir, ABH`s field writer from Germany, hunted a