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A RT I C L E S
Implications of Deep-Marine Miocene Deposits on the
Evolution of the North Patagonian Andes
Alfonso Encinas,1,* Patricio A. Zambrano,1 Kenneth L. Finger,2
Victor Valencia,3 Luis A. Buatois,4 and Paul Duhart5
1. Departamento de Ciencias de la Tierra, Universidad de Concepción, Casilla 160-C, Concepción, Chile;
2. University of California Museum of Paleontology, Valley Life Sciences Building 1101, Berkeley, California
94720, USA; 3. School of Earth and Environmental Sciences, Washington State University, Pullman,
Washington 99164, USA; 4. Department of Geological Sciences, University of Saskatchewan,
114 Science Place, Saskatoon, Saskatchewan S7N 5E2, Canada; 5. Servicio Nacional de
Geologı́a y Minerı́a, Oficina Técnica Puerto Varas, Casilla 613, Puerto Varas, Chile
ABSTRACT
The North Patagonian Andes (38⬚–43⬚30 S) present topographic and geologic characteristics distinct from those of the
Central Andes to the north, including a lower altitude, a predominance of plutonic rocks, and the presence of a major
N-S-trending, dextral strike-slip, intra-arc discontinuity known as the Liquiñe-Ofqui Fault Zone (LOFZ). The timing
for the uplift of the North Patagonian Andes remains poorly constrained principally because high exhumation rates
resulted in the erosion of most of its volcano-sedimentary cover during the late Cenozoic. The strongly deformed
Cenozoic marine deposits ascribed to the Ayacara Formation, which crop out on the western flank of this range at
∼42⬚S, can provide for a better understanding of the recent geologic evolution of this range. To discern this formation’s
age and sedimentary environment, we have integrated sedimentology, ichnology, foraminiferal paleontology, and
geochronology. Our results indicate that this unit was deposited in a deep-marine environment during the EarlyMiddle Miocene. Correlation of the Ayacara Formation with coeval deep-marine deposits in the forearc to the west
indicate that deposition was caused by a major regional event of tectonic subsidence probably related to subduction
erosion. An Early-Middle Miocene age for the Ayacara Formation is a reliable maximum age for the deformation and
uplift of the western flank of the North Patagonian Andes as well as for the commencement of transpressional
deformation associated with the LOFZ.
Online enhancements: appendix, supplementary tables.
Introduction
the lower plates that characterizes the “Chilean
subduction mode.” They ascribed such a strong
coupling to a high convergence rate and a relatively
young and buoyant oceanic plate. More than 30 yr
of research, however, has shown that their model
is too simplistic, and several issues remain highly
debated. First, the uplift of the Andean Cordillera
has been ascribed to various causes, such as (a) the
convergence rate between the oceanic and continental plates (e.g., Pardo-Casas and Molnar 1987);
(b) the absolute westward motion of South America
(e.g., Silver et al. 1998); (c) the amount of sediment
along the plate interface, which determines shear
The Andean Cordillera, the longest and highest
mountain belt formed in a convergent setting, extends for ∼4000 km along the Chilean margin. The
origin of this range is related to the subduction of
the oceanic Nazca (formerly Farallon) plate beneath
South America, at least since the Jurassic (Mpodozis and Ramos 1989). Uyeda and Kanamori (1979)
related the formation of the Andean range to the
strong mechanical coupling between the upper and
Manuscript received December 12, 2011; accepted December
4, 2012.
* Author for correspondence; e-mail: [email protected].
[The Journal of Geology, 2013, volume 121, p. 215–238] 䉷 2013 by The University of Chicago.
All rights reserved. 0022-1376/2013/12103-0001$15.00. DOI: 10.1086/669976
215
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A. ENCINAS ET AL.
stress in the subduction zone and is ultimately controlled by climate (Lamb and Davis 2003); (d) the
age of the subducting plate (Yañez and Cembrano
2004); (e) a combination of the overriding velocity
of the South American plate, crustal thickness, and
friction coefficient in the subduction channel
(Sobolev and Babeiko 2005); and (f) the subduction
of horizontal slab segments (Martinod et al. 2010).
Second, the long-lived Andean margin has not been
in a permanent state of compression, as envisaged
by Uyeda and Kanamori’s model. On the contrary,
it appears that tectonic shortening and Andean uplift occurred principally during three pulses in the
upper Cretaceous, Eocene, and Neogene, whereas
other periods are characterized by moderate shortening or extension (Martinod et al. 2010 and references therein). Third, despite rather uniform
convergence conditions, the Andean margin is
characterized by significant along-strike segmentation (e.g., Mpodozis and Ramos 1989; Martinod
et al. 2010; Orts et al. 2012). Maximum mean elevations, width, and crustal shortening in the Central Andes are much larger than those in the southern and northern parts (e.g., Martinod et al. 2010
and references therein). Segmentation is also evident in the margin’s geologic record, as tectonic
phases that caused Andean growth did not occur
uniformly along the whole mountain chain (e.g.,
Jordan et al. 2001; Martinod et al. 2010 and references therein). Fourth, although the Andean Cordillera is considered to have formed by crustal
shortening through the advance of predominantly
east-verging fold and thrust belts (Farı́as et al. 2010
and references therein), some areas are characterized by important margin-parallel strike-slip systems whose origin is debated (e.g., Cembrano et al.
2002; Thomson 2002).
The North Patagonian Andes (38⬚–43⬚30S) are a
poorly studied segment of the Andean Cordillera
that present important differences compared with
the Central Andes to the north. South of ∼38⬚S, the
broad and high (14000 m) Central Andes narrow
considerably, and maximum elevation descends to
∼2000 m (Hervé 1994). The North Patagonian Andes have a crustal thickness of only ∼40 km (Lüth
et al. 2003), in contrast to the 70 km for the Central
Andes (Allmendinger et al. 1997). They are mainly
composed of plutonic rocks of the Patagonian Batholith (Adriasola et al. 2006), show little deformation
(Hervé 1994; Orts et al. 2012), and have a different
timing of the main contractional phases than the
Central Andes (Garcı́a Morabito et al. 2011and references therein).
A distinctive tectonic feature of the North Patagonian Andes is a ∼1000-km-long, N-S-trending,
dextral strike-slip fault system referred to as the
Liquiñe-Ofqui Fault Zone (LOFZ; Hervé 1994). This
fault system accommodates part of the marginparallel component of oblique subduction. However,
the long-term nature, timing, and causes of lateral
motion along this important fault system remain
poorly understood. Also distinctive of this segment
of the Andean margin is the subduction of the Chile
Rise spreading center that defines the Chile triple
junction between the Nazca, South American, and
Antarctic plates. Subduction of the Chile Rise beneath the southernmost tip of the South American
plate began at ∼14 Ma and migrated north to its
present location at ∼46⬚S (Adriasola et al. 2006). Finally, an important singularity of the geologic evolution of the North Patagonian Andes is the occurrence of late Cenozoic marine deposits across the
entire forearc and in the western and eastern flanks
of this range, indicating that uplift of this part of the
Andes occurred rather recently (Levi et al. 1966;
Ramos 1982; Encinas et al. 2012b).
Geologic studies of the North Patagonian Andes
have been limited by a paucity of Meso-Cenozoic
strata. High exhumation rates during the late Cenozoic resulted in extensive erosion of the volcanosedimentary cover (Adriasola et al. 2006), and what
remains is mostly obscured by thick vegetation. As
a consequence, there is little stratigraphic control
to constrain the timing and onset of deformation.
However, study of Cenozoic marine strata exposed
on the western flank of this range at ∼42⬚S (figs. 1,
2) can enable us to better understand the recent
geologic evolution of the North Patagonian Andes.
These marine strata, referred to as the Ayacara Formation (Levi et al. 1966), are strongly deformed and
constitute a rhythmic succession of sandstone and
siltstone. The age of this unit has been ascribed to
the Eocene-Miocene (Levi et al. 1966; Rojas et al.
1994), Late Eocene–Late Oligocene (Bourdillón, in
SERNAGEOMIN 1995), and Miocene (Martı́nezPardo 1965). Another unresolved issue is its possible correlation with the Neogene marine strata of
the Chilean forearc, which intrigued the first geologists who studied this area (e.g., Katz 1965).
Hervé (1994), Rojas et al. (1994), and Hervé et al.
(1995) speculated that the Ayacara Formation and
correlative units to the south formed in short-lived
pull-apart basins related to the dextral motion of
the nearby LOFZ. As noted by Cembrano et al.
(2002), this interpretation is inconclusive, as there
is little structural and chronological evidence to
support it.
In this study, we investigate the formation and
subsequent destruction of a forearc basin at a particular segment of an ocean-continent convergent
Figure 1. Geologic map of the study area. Insets show the location of the studied sections detailed in figure 2. (After
Ramos 1982; SEGEMAR 1995; Giacosa et al. 2005; Duhart 2008.) LOFZ p Liquiñe-Ofqui Fault Zone.
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A. ENCINAS ET AL.
Figure 2. Detailed geologic maps showing sections from figure 3 (arrows) as well as the strike and dip of beds and
inverse fault at Fiordo Pichicolo depicted in figures A3 and A4, available in the online edition or from the Journal
of Geology office. LOFZ p Liquiñe-Ofqui Fault Zone.
margin that has been influenced by strike-slip tectonics and the subduction of an active mid-oceanic
ridge. For that purpose, we focus on the marine
deposits of the Ayacara Formation that crop out on
the western flank of the North Patagonian Andes.
We integrate sedimentology, ichnology, foraminiferal micropaleontology, and laser ablation multicollector inductively coupled plasma mass spectrometry (LA-MC-ICPMS) U-Pb geochronology on
detrital zircons to discern the age and sedimentary
environment of this unit. Our findings have important implications for the geologic evolution of
this segment of the Andean Cordillera, as they (1)
indicate the onset of a major event of subsidence
that affected the entire forearc and at least the western flank of the North Patagonian Andes and (2)
provide a reliable maximum age for the uplift and
deformation of the western flank of the range. On
the basis of our findings and previous work in this
region, we delve into the causes of forearc subsidence during the Early-Middle Miocene and the
subsequent onset of uplift and transpressional deformation of this segment of the Andean Cordillera.
Geologic Setting
The study area is located between the Estuario
Reloncavı́ and the Peninsula Huequi (∼42⬚S) on the
western flank of the North Patagonian Andes (figs.
1, 2). The three N-S-trending physiographic units
in this region (fig. 1) are, from west to east, the
Coastal Cordillera, the Intermediate Depression,
and the Main Andean Cordillera. The Coastal Cordillera is a subdued mountain range located in the
western part of Chiloé Island (Duhart and Adriasola
2008). The Intermediate Depression (also referred
to as the Longitudinal Depression or the Central
Valley) is a low-lying area between the Coastal Cordillera and the Main Andean Cordillera that includes eastern Chiloé Island and the adjacent floor
of the Golfo de Ancud (Duhart and Adriasola 2008).
The Main Andean Cordillera extends from the
mainland coast to eastern Argentina and includes
the highest peaks and active volcanoes (Duhart and
Adriasola 2008).
The Coastal Cordillera (fig. 1) has a PaleozoicTriassic metamorphic basement interpreted as a paleoaccretionary wedge (Duhart and Adriasola 2008).
Locally, Eocene dacitic dikes and sills, as well as
granodioritic stocks, intrude the basement (Duhart
and Adriasola 2008). Coal-bearing rocks of the Caleta Chonos Formation (Eocene?-Oligocene?) occur
in a small area of northwestern Chiloe (Antinao et
al. 2000). Late Oligocene–Early Miocene volcanic
and subvolcanic rocks of the Complejo Volcánico
de Ancud and equivalent units are mostly exposed
in northwestern Chiloé (Antinao et al. 2000).
Younger Neogene marine strata of the Lacuı́ Formation and equivalent units overlie the metamorphic basement or the volcanic rocks of the Com-
Journal of Geology
PATA G O N I A N D E E P - M A R I N E M I O C E N E D E P O S I T S
plejo Volcánico de Ancud (Antinao et al. 2000;
Duhart and Adriasola 2008; Encinas et al. 2009).
The age of the Lacuı́ Formation and correlative
units is controversial. Molluscan biostratigraphy
and Sr isotopes indicate a latest Oligocene–late
Early Miocene age (Finger et al. 2007; Nielsen and
Glodny 2009). Samples from different sections
yielded three different assemblages of planktic foraminifera, indicative of the Burdigalian-Serravallian (Antinao et al. 2000), the Tortonian (Finger et
al. 2007), and the Zanclean (Finger et al. 2007). According to these studies, this unit ranges from Early
Miocene to Early Pliocene, but the younger markers reported in Finger et al. (2007) were misidentified, and most of the assemblages are probably
Early Miocene. Benthic foraminifera indicate
lower-bathyal depositional depths for the Lacuı́ Formation (Finger et al. 2007). Shallow-marine Pliocene? strata of the overlying Caleta Godoy Formation occur in some parts of northern Chiloé
Island (Antinao et al. 2000). Pleistocene deposits,
mostly glacial, and Holocene continental and marine sedimentary beds constitute the younger successions in the area.
The geology of the Longitudinal Depression (fig.
1) in this area is poorly known because most of this
physiographic feature is submerged. Subsurface
data from a 4010-m well drilled by ENAP (the Chilean National Oil Company) at Puerto Montt (Katz
1965; Elgueta et al. 2000) transected a succession
comprising (1) a lower, 1480-m-thick interval of interbedded tuff and volcaniclastic breccia; (2) 310 m
of Miocene marine rocks that probably correlate
with the Lacuı́ Formation; (3) 950 m of Pliocene?
continental? or marine? sedimentary rocks; and (4)
1300 m of Pleistocene continental and marine deposits mostly of glacial origin. Miocene? marine
strata and Pleistocene marine and continental deposits also occur in some areas along the coast of
eastern Chiloe and on the nearby islands (Quiroz
et al. 2004; Duhart and Adriasola 2008).
The oldest rocks of the Main Andean Cordillera
(fig. 1) are Middle-Late Paleozoic metamorphics in
the western part (Duhart 2008). Mesozoic-Cenozoic plutonics of the North Patagonian Batholith
(Pankhurst et al. 1992) are the most extensive rocks
in the Main Andean Cordillera. Jurassic, Cretaceous, Paleogene, and Neogene volcanic, volcaniclastic, and sedimentary rocks occur on the eastern
and western flanks (Giacosa et al. 2005; Duhart
2008). In Chile, Miocene marine sedimentary rocks
of the Ayacara and La Cascada Formations (Levi et
al. 1966; Thiele et al. 1978; Encinas et al. 2012a)
occur along the western and eastern parts, respectively. Miocene marine deposits of the Rio Foyel
219
Formation occur in westernmost Argentina (Bertels
1980; Ramos 1982; Encinas et al. 2012a). Pleistocene and Holocene volcanic, volcaniclastic, and
glacial deposits constitute the youngest successions in the North Patagonian Andes.
The LOFZ extends for ∼1000 km along the North
Patagonian Andes (figs. 1, 2) and decouples the
south-central Chilean forearc, which behaves as a
northward-gliding sliver, according to Forsythe and
Nelson (1985). The long-term nature and timing of
lateral motion along this important fault system,
however, remain poorly understood (Thomson
2002). Structural analyses, Ar-Ar geochronology,
and fission-track thermochronology along the
Northern Patagonian Andes (41⬚–46⬚S) provide
solid evidence for right-lateral transpression along
the LOFZ during the Late Miocene to Pliocene (e.g.,
Cembrano et al. 2000, 2002; Thomson 2002;
Adriasola et al. 2006). However, Cembrano et al.
(2000) document a 1-km-wide mylonitic zone that
shows microstructural evidence of sinistral reverse-displacement and is crosscut by a 100-Ma
(Ⳳ2 Ma) undeformed porphyritic dike at Liquiñe
(39⬚S). The LOFZ appears to be linked to a zone of
magmatic weakening of the crust, as suggested by
alignment of its strike with Pliocene-Recent volcanic centers (Hervé 1994; Glodny et al. 2008).
Previous Studies of the Ayacara Formation
The Ayacara Formation (Levi et al. 1966; figs. 1, 2)
is a succession of turbidites composed of sandstone,
siltstone, conglomerates, and minor tuffs between
the Estuario Reloncavı́ and Chaitén, with the most
complete sections at Caleta Ayacara and Isla El
Manzano (figs. 1, 2). Its scarce fossil fauna includes
solitary corals (Solano 1978), echinoderms (Alarcón
1995), crustaceans (Feldmann et al. 2010), foraminifers, and silicoflagellates (Levi et al. 1966;
SERNAGEOMIN 1995). Its strata are intruded by
igneous dikes (Levi 1966; Duhart 2008).
Levi et al. (1966) assigned an Eocene-Miocene
age to the Ayacara Formation on the basis of Martı́nez-Pardo’s study of foraminifers and silicoflagellates. Martı́nez-Pardo (1965) ascribed a Middle
Miocene age to a foraminifer sample from Isla El
Manzano. This age was supported by Fuenzalida
(1979), who identified the planktic foraminifer
Orbulina universa in the Ayacara Formation at
Peninsula Huequi. Bourdillón (1994) criticized
these studies because the foraminifera were identified only from petrographic thin sections, a
methodology of limited accuracy. He later extracted foraminifers from several Ayacara Formation samples that were indicative of a Middle
Figure 3. Representative sections from Isla El Manzano (IMAN), Caleta Ayacara (CAYA), and Isla Ica (IICA; see fig.
3 for location). LM p lower member, UM p upper member. Diagonal lines in the scales represent long covered
intervals. Grain size: (a) clay, (b) silt, (c) very fine sand, (d) fine sand, (e) medium sand, (f) coarse sand, (g) very coarse
sand, (h) granule, (i) pebble, (j) cobble, (k) boulder. LT p lithology: (1) clast-supported conglomerate, (2) matrixsupported conglomerate, (3) breccia, (4) sandstone, (5) siltstone, (6) tuff, (7) andesitic dike, (8) covered interval. SS p
sedimentary structures: (9) trough cross-bedding, (10) planar cross-bedding, (11) hummocky cross stratification, (12)
Journal of Geology
Eocene–Late Oligocene age (in SERNAGEOMIN
1995). Radiometric 40Ar/39Ar dating of plagioclase
from an Ayacara Formation tuff near the top of the
Caleta Ayacara succession, however, yielded an
age of 16.5 Ⳳ 0.5 Ma (F. Munizaga, written comm.
in Rojas et al. 1994).
221
units are obscured by dense vegetation. However,
the intense deformation of this unit in some localities, as well as the occurrence of exhumed Miocene granites in this region, suggest that some of
these contacts are inverse faults.
Sedimentary Facies of the Lower Member
Sampling Sites
Because of the thick vegetation cover in the study
area, Ayacara Formation exposures are limited to
coastal cliffs and a few roadcuts and river valleys.
The formation underlies Pleistocene? deposits
(Levi et al. 1966), but its basal contact has not been
observed. We performed detailed sedimentologic
and ichnologic analyses and measured stratigraphic
thicknesses at Isla El Manzano, Caleta Ayacara, and
Isla Ica (figs. 2, 3), where the unit is thickest and
least weathered. In addition, we collected rock samples for foraminiferal analysis and U-Pb zircon
geochronology.
This unit was divided into two members by Levi
et al. (1966). The lower member, 180 m thick, crops
out only at the Caleta Ayacara section, whereas the
upper member, up to 570 m thick, occurs in all of
the studied localities (fig. 3). The transition between the two members at Caleta Ayacara is
covered.
We also measured bedding attitudes in different
outcrops (fig. 2). The average strike and dip of the
beds are 335⬚/75⬚ SW in the lower member and
325⬚/75⬚ SW in the upper member at Caleta Ayacara, 335⬚/85⬚ SW at Isla Ica (fig. A1, available in
the online edition or from the Journal of Geology
office), and 35⬚/40⬚–70⬚ NW at Isla El Manzano.
Strata of the coastal outcrops of Fiordo Pichicolo
average 360⬚/66⬚ W but become vertical to the east
(fig. A2, available in the online edition or from the
Journal of Geology office). In a newly opened roadcut that runs parallel to this fjord, we measured a
270⬚/75⬚ W reverse fault but did not observe slickenslides. The fault has associated tight folds and a
30-cm-thick tectonic breccia (figs. A3, A4, available
in the online edition or from the Journal of Geology
office).
Ayacara Formation contacts with other geologic
Matrix-Supported Conglomerate (Cgl1). This facies
consists of sharp- to erosively based massive,
matrix-supported conglomerate (fig. 4A) that forms
meters-thick successions. It is poorly sorted and
contains subangular to subrounded clasts composed primarily of volcanic rocks (andesite with
aphanitic and traquitic textures, tuff, and minor
dacite) and secondarily of plutonic rocks (diorite
and tonalite) and sandstone. Clasts show random
orientations and range in size from millimeters to
several decimeters. This facies locally grades upward into the clast-supported conglomerate of facies Cgl2.
This facies is interpreted as debris-flow deposits
on the basis of their poor sorting, randomly oriented floating clasts, and abundance of matrix (e.g.,
Nemec and Steel 1984; Hwang and Chough 1990).
Clast-Supported Conglomerate (Cgl2). This facies
consists of slightly erosive to, more rarely, sharpbased clast-supported conglomerate that generally
has a lot of sandy matrix. Beds range in thickness
from decimeters to, more commonly, meters. Conglomerate is generally well to moderately sorted.
Clasts range in size from a few millimeters to several centimeters, albeit outsized boulders up to 80
cm occur locally. Clasts are subangular to rounded
and compositionally similar to those of facies Cgl1.
Some beds show abundant rip-up sandstone clasts
up to 40 cm across. Clasts are generally arranged
parallel to bedding, and some are imbricated. Beds
are generally massive or show crude stratification.
However, some strata show normal or inverse gradation and, more rarely, planar cross-bedding (fig.
4B). Conglomerates in this facies sometimes grade
into sandstones of facies Ls.
Clasts orientated parallel to bedding or imbricated in this facies indicate collision during transport (Hwang and Chough 1990). The presence of
asymmetrical ripples, (13) symmetrical ripples, (14) parallel lamination, (15) massive, (16) convolute lamination, (17)
flames, (18) water-escape structures, (19) dish structures, (20) load structures, (21) flute casts, (22) ball and pillow
structures, (23) slumps, (24) imbricated clasts, (25) floating clasts, (26) carbonated concretions, (27) pyrite concretions,
(28) rip-up clasts. F p fossils: (29) foraminifers, (30) bivalves, (31) gastropods, (32) scaphopods, (33) brachiopods, (34)
solitary corals, (35) echinoderms. TF p trace fossils: (36) Thalassinoides, (37) Ophiomorpha, (38) radially branching
burrows, (39) Planolites, (40) Schaubcylindrichnus, (41) Paleophycus, (42) Taenidium, (43) Zoophycos, (44) Chondrites,
(45) Phycosiphon, (46) undetermined burrows.
Figure 4. Characteristic facies and trace fossils of the lower member of the Ayacara Formation at Caleta Ayacara
section. A, Matrix-supported, poorly sorted conglomerate (facies Cgl1). B, Clast-supported conglomerate (facies Cgl2).
The bed indicated by the hammer shows planar cross-bedding. The upper bed is sandstone. C, Coarsening-upward
parasequence formed by a basal sandy-siltstone interval that grades into sandstone to the top (facies Ls). The base
(flooding surface) is indicated by the arrow. D, Hummocky cross stratified (HCS) sandstone (facies Ls). Note the lowangle curved intersections of stratification characteristic of HCS, indicated by the pen. E, Sandstone with symmetrical
wave ripples (facies Ls). F, Undetermined radially branching burrows in sandy siltstones from the basal interval of
the parasequence shown in C.
Journal of Geology
grading illustrates that clasts within the flow were
free to move relative to each other, suggesting little
or no cohesion (Van der Straaten 1990). Planar
cross-bedding represents 2-D dune migration.
These sedimentary features indicate transport by
fluidal flows. The common abundance of sand matrix and outsized clasts suggests that transport by
hyperconcentrated flows was the dominant process
(Iverson 1997).
Laminated Sandstone (Ls). This facies consists
principally of laminated sandstone interbedded
with siltstone or conglomerate forming two different associations. Facies association Ls1 consists of
meters-thick, coarsening-upward successions characterized by a basal sandy-siltstone interval that
grades upward into sandstone (fig. 4C). Siltstone is
dark gray and in sharp contact with the underlying
rock. Sandstone is fine grained to, more commonly,
medium to coarse grained and usually shows hummocky cross stratification (fig. 4D) and scarce intercalations of conglomerate a few centimeters
thick. Facies association Ls2 comprises a mediumto very coarse-grained sandstone interbedded with
conglomerate of facies Cgl2. Some beds have an
erosional basal contact overlain by a conglomerate
that grades into sandstone. Strata are typically tabular and decimeters to a few meters thick. Some
sandstone beds show symmetrical ripples (fig. 4E),
while others display trough or planar cross-bedding.
Sandstone is composed primarily of quartz, andesite lithics, and plagioclase and secondarily of clinopiroxene, although some beds consist of shards,
pumice, and quartz.
Sandstone in this facies shows sedimentary
structures that indicate wave reworking. Facies association Ls1 is interpreted as shallowing-upward
parasequences, where basal siltstones represent
shelf deposition and hummocky cross stratified
sandstone indicates reworking by storm waves in
the lower shoreface (Clifton 2006). Thin intercalations of conglomerates are interpreted as storm
deposits. Laminated sandstone in facies association
Ls2 reflects wave influence. Sandstones showing
symmetrical ripples are typical of the lower shoreface, whereas those with cross-bedding indicate deposition in the upper shoreface (Clifton 2006).
The Ayacara Formation’s lower member presents
an overall fining- and thinning-upward succession,
where facies Cgl1 (matrix-supported conglomerates) is present only in the lower part. Facies Cgl2
(clast-supported conglomerates) occurs along the
whole interval but is more abundant and thicker
in the lower part, and facies Ls (laminated sandstone and minor siltstone) is present only in the
upper half of the succession.
223
Sedimentary Facies of the Upper Member
Rhythmically Interbedded Sandstone and Siltstone
(T). This facies consists of fine- to very coarse-
grained sandstone rhythmically interbedded with
siltstone (fig. 5A). Bed thickness is variable, but
sandstone beds are typically decimeters to meters
thick, whereas siltstone intervals are typically centimeters to decimeters thick. They form partial
Bouma cycles where sandstone predominates and
is locally amalgamated. They typically show Tbc
divisions (parallel-convolute or, more rarely, ripple
cross-lamination; fig. 5B), but Tab (structurelessparallel-laminated), Ta, and Tb are also common.
Sandstone has sharp and, more rarely, erosive basal
contacts and flute casts (fig. 5C) and occasionally
contains ripped-up mudstone clasts up to some
decimeters in size. Sandstone beds commonly present convolute lamination, locally overturned, in
intervals up to 1 m thick. Fine-grained (Tde) intervals consist of siltstone and, more rarely, tuff. They
are generally azoic, but some of the tuffaceous siltstone contains abundant foraminifers. Flames, load
casts, and, more rarely, ball and pillow structures
locally occur at the contact between sandstone and
siltstone beds. Sandstone is primarily composed of
plagioclase, quartz, and volcanic lithics and secondarily of shards, pumice, and igneous lithics. Interbedded sandstone and sitstone (T) and conglomerate (Cgu) are the most common facies in the
upper member of the Ayacara Formation.
The rhythmically interbedded sandstone- and
siltstone-forming Bouma cycles indicate that this
facies represents classic turbidites (Bouma 1962;
Posamentier and Walker 2006). The predominance
of basal Bouma divisions (Ta, Tb), as well as the
coarse grain size and thick bedding, suggests proximal turbidites. The common presence of convolute lamination, as well as the occurrence of
flames, load casts, and ball and pillow structures,
indicates rapid sediment deposition that causes
trapping of pore fluid and deformation of primary
structures (Posamentier and Walker 2006). The
presence of rip-up clasts and flute casts in some
beds are evidence of turbulent erosion of the substrate immediately prior to deposition.
Conglomerate (Cgu). Individual conglomerate
beds are typically decimeters to meters thick and
are usually interbedded with turbidites, sandstones, or siltstones. Basal contacts are sharp or
slightly erosional and sometimes show load and
flame structures (fig. 5D), pseudonodules, descendent clastic dikes, or disruption of underlying siltstone beds. Conglomerate is clast supported, is
moderately to well sorted, and contains subangular
Figure 5. Characteristic facies and trace fossils of the upper member of the Ayacara Formation. A, Rhythmically
interbedded sandstone and siltstone (facies T) at Isla Ica. A human for scale is encircled. B, Sandstone showing partial
Tbc (parallel-convolute lamination) Bouma cycle (facies T) at Caleta Ayacara. C, Base of turbidite with flute casts
(facies T) at Isla Ica. Paleocurrent direction (157⬚) is indicated by the arrow. D, Flame formed by tuffaceous siltstone
injecting into granule conglomerate (facies Cgu) at Isla El Manzano section. E, Slumped tuff beds (facies Slp) at Caleta
Ayacara section. F, Chondrites isp. in very fine-grained sandstone at Isla Ica section. Note the typical branching of
this trace fossil.
Journal of Geology
to rounded clasts, typically a few centimeters in
size. Clasts are mainly aphanitic and traquitic andesite and minor sedimentary and metamorphic
lithics. They present a sandy matrix composed of
angular granules of plagioclase and quartz, some of
them showing embayments. Conglomerate occasionally contains rip-up clasts up to a few decimeters in size that consist of fragments of stratified
siltstones, tuffs, or interbedded siltstone and sandstone, some of which display folding. Some beds,
however, are composed entirely of angular rip-up
clasts. On the basis of Walker’s (1975) classification, we distinguished three different types of conglomerates: the Cgu1 type is normally graded with
subtle stratification, generally grading upward into
pebbly sandstone and rarely showing planar crossbedding; the Cgu2 type is normally graded and not
stratified; and the Cgu3 type is inversely to normally graded. Conglomerate Cgu2 is predominant,
Cgu1 is less common, and Cgu3 is rare. Normal
grading in Cgu1 and Cgu2 is generally very subtle
and may be absent.
The presence of grading and, more rarely, crossbedding in this facies suggest transport by turbulent
flows in which grains move freely relative to one
another. Conglomerates in deep-marine successions that show similar facies to those described
above are interpreted as having been deposited by
turbidity currents (Posamentier and Walker 2006).
The presence of rip-up clasts is evidence of turbulent erosion of the substrate immediately prior
to deposition, whereas the occurrence of disruptive
sedimentary structures, such as flames and load
casts, indicates rapid deposition.
Structureless Sandstone (Ss). Structureless sandstones are medium- to very coarse-grained, usually
massive sandstones, although they may locally display zones of indistinct parallel lamination, and
some beds show dish structures. Structureless
sandstone successions are decimeters to meters
thick and locally form amalgamated stacks that
have erosional contacts or rip-up clasts occasionally folded up to 1 m in size between individual
beds. Mineralogical composition of this facies is
similar to that of sandstones in facies T.
Structureless sandstones are commonly interpreted as deposited by turbidity currents on the basis of the common association of this facies with
classic turbidites (Posamentier and Walker 2006).
The presence of dish structures that indicate initial
deposition of a fluid-rich sediment-water mixture
also suggests this mode of deposition.
Siltstone and Very Fine-Grained Sandstone (Sil).
This facies consists of dark gray to brownish siltstone and very fine-grained sandstone forming in-
225
tervals up to several meters thick. Locally it shows
thin intercalations of coarse-grained sandstone and
granule conglomerate. Bedding thickness ranges
from centimeters to decimeters. Beds are massive
or show poorly defined lamination. Fossils of bivalves, gastropods, scaphopods, echinoderms, brachiopods, and solitary corals are, with few exceptions, present only in this facies of the Ayacara
Formation. Foraminifers, mostly planktic, are
abundant, and pyrite is common.
This facies is interpreted as having been deposited by distal turbidity currents and hemipelagic
fallout. Interbedded sandstone and conglomerate
are coarser-grained turbidites. Thick, fine-grained
intervals within turbiditic successions are usually
interpreted as condensed sections deposited at extremely low sedimentation rates during maximum
regional transgression (Weimer and Slatt 2006). The
abundant and diverse biota, the high ratio of planktic to benthic foraminifers, the high content of organic matter suggested by the dark color of this
facies, and the presence of pyrite are also characteristic features of condensed sections (Weimer and
Slatt 2006).
Tuff (Tf). This facies consists principally of ash
and secondarily of lapilli tuff and is commonly interbedded with sandstone and conglomerate. Beds
typically show a green color and are commonly
decimeters thick, but at the Caleta Ayacara section
there are intervals up to 18 m thick. Tuff consists
mostly of shards and subordinate pumice, quartz,
and plagioclase (fig. A5A, A5B, available in the online edition or from the Journal of Geology office).
Shards are generally transformed to zeolites. This
facies usually shows a massive aspect, but parallel
lamination, convolute bedding, and, more rarely,
slumps are present in some beds. Siltstone rip-up
clasts up to 80 cm across, some of which are deformed, are present locally.
This facies was initially formed by explosive volcanism (ashfalls and pyroclastic flows), as indicated
by the abundance of shards and pumice clasts. Finegrained volcanic material entered water and was
deposited by settling. Coarser-grained pyroclastic
flows were transformed into water-supported turbidity flows that displaced them farther and deeper
into the marine basin, as suggested by the presence
of parallel and convolute laminations (see Cas and
Wright 1991).
Slumped Sandstone and Siltstone (Slp). This facies
only rarely occurs and consists of folded interbedded sandstone and siltstone forming intervals decimeters to a few meters thick. These beds are underlain and overlain by flat-bedded successions (fig.
5E).
226
A. ENCINAS ET AL.
This facies was formed by sliding and folding of
unconsolidated turbidites. They were previously
deposited in the basin and subsequently displaced
downslope.
Trace Fossils
Trace fossils in the lower member of the Ayacara
Formation are indeterminate radially branching
burrows (fig. 4F) present only in siltstone and very
fine-grained sandstone. Bioturbation in the upper
member is scarce and present only in discrete intervals (!5% of beds are bioturbated). Two distinct
trace-fossil assemblages can be differentiated.
Trace-fossil assemblage 1 occurs in siltstone and,
more rarely, very fine-grained sandstone of facies T
and Sil. It is dominated by Chondrites isp. (fig. 5E),
with Phycosiphon incertum and Planolites montanus accessory elements and Zoophycos isp.,
Schaubcylindrichnus freyi, and Taenidium isp. rare
components. The bioturbation index (BI scale by
Taylor and Goldring 1993) is typically low (1–2),
although it reaches 5 in some intervals, which are
also characterized by an increase in ichnodiversity.
This assemblage is ascribed to the Zoophycos ichnofacies, characterized by low oxygen levels associated with high organic detritus in quiet-water settings, and is typical of outer shelf to slope settings
(Frey and Pemberton 1984). In the case of the
Ayacara assemblage, it essentially characterizes interturbidite intervals of colonization.
Trace-fossil assemblage 2 occurs in medium- to
coarse-grained sandstone of facies T and Ss and,
more rarely, in tuff (facies Tf). It is dominated by
Thalassinoides suevicus and Ophiomorpha isp.,
each of which commonly forms a monospecific
suite characterized by extensive horizontal networks. Palaeophycus tubularis is locally present.
Although typically low, the bioturbation index is
somewhat variable, reaching BI 4 at the network
levels. This trace-fossil assemblage characterizes
the Skolithos ichnofacies, typical of high-energy
shallow-marine environments but also widely reported in deepwater settings, and it reflects local
environmental conditions, such as high energy, a
sandy substrate, high levels of oxygenation, and an
abundance of suspended organic particles (Buatois
and López-Angriman 1992). The low bioturbation
index of the deposits, the low trace-fossil diversity
of both associations, and the predominance of morphologically simple structures probably reflect
short colonization windows in a stressful environment characterized by rapid rates of volcaniclas-
tic sedimentation (see also Mángano and Buatois
1997).
Foraminiferal Biostratigraphy and Paleodepths
To determine the age and paleobathymetry of the
Ayacara Formation, we analyzed nine samples from
the Caleta Ayacara and Isla El Manzano sections.
Details about the extraction method are provided
in the appendix, available in the online edition or
from the Journal of Geology office. All of the isolated specimens show some degree of diagenetic
alteration that masks diagnostic features. Thus,
species identifications (table S1, available in the
online edition or from the Journal of Geology office)
were based mostly on experience and intuition
stemming from familiarity with the Miocene fauna
of south-central Chile.
Planktic foraminifera are generally more abundant than benthic specimens in the analyzed samples, reaching P : B ratios of 90–100 in some beds
of the Caleta Ayacara section, where planktics are
so profuse that they constitute as much as 75% of
the rock volume. Only five samples, all from the
Caleta Ayacara section, yielded planktic specimens
that could be identified at the species level (tables
S1, S2, available in the online edition or from the
Journal of Geology office). Samples CAYA 17, 18,
19, 20, and 22 yielded Paragloborotalia mayeri, a
Late Oligocene–Middle Miocene species with a
zonal range of P22–N14. CAYA 22 also yielded Globigerinoides quadrilobatus, and the two species
have a concurrent range of N6–N14 (BurdigalianSerravallian). Finally, the occurrence of Globigerinoides sicanus in CAYA 17 restricts the sample to
zones N8–N9 (Langhian, Middle Miocene). This determination narrows the range for four stratigraphically higher samples to N8–N14 (Langhian-Serravallian, Middle Miocene).
Benthic foraminifera were present in four samples from the Caleta Ayacara section and in three
samples from the Isla El Manzano sections (table
S1), but all of the benthic assemblages were weak
(i.e., consisting of relatively few specimens). Depositional paleodepths are derived from the upperdepth limits (UDLs) of benthic foraminifera in the
Chilean Miocene, which are based on the depth
distributions of similar taxa currently living along
the Pacific margin of central South America, as recorded by Bandy and Rodolfo (1964), Ingle et al.
(1980), and Resig (1981). Most of the studied assemblages are mixed-depth associations of shelf and
slope species, indicating downslope displacement
(table S1). Five samples (CAYA 10, 18, 22; IMAN
Journal of Geology
16, 30) yielded foraminifera indicative of uppermiddle bathyal (500–1500 m) depths—Cyclammina incisa (Stache), Gyroidinoides nipponica
(Ishizaki), Hansenisca sp., and Neouvigerina hispida (Schwager)—some of which are included in the
cosmopolitan deepwater fauna of Van Morkhoven
et al. (1986). One sample (CAYA 10) also had
Siphonodosaria sagrinensis (Bagg), which occurs no
shallower than lower-middle bathyal (1500–2000
m) depths. In addition, a fragment possibly of that
species was found in IMAN 30. Considering the
poor preservation and weak benthic assemblages,
there can be little doubt that these occasional specimens represent a much more diverse fauna that
most likely was deposited below 1500 m.
U-Pb Geochronology
Five samples, each consisting of ∼10 kg of mediumto coarse-grained sandstones, were selected for UPb zircon geochronology, which was conducted by
LA-MC-ICPMS at the Arizona LaserChron Center
following the techniques described by Gehrels et
al. (2008) and detailed in the appendix. Two of the
five samples were from the Isla El Manzano section,
and three samples were from the Caleta Ayacara
section (table S3; fig. 6). Four samples correspond
to the upper member of the Ayacara Formation
(CAYA 14, CAYA 24, IMAN 14, and IMAN 20), and
one sample (CAYA 4) belongs to the lower member
of this unit. Crystals in these samples are predominantly prismatic and have a clear pinkish color
except for a few reddish, subrounded grains in sample CAYA 4. The U/Th ratio is !7 in the grains
from the Isla El Manzano and Caleta Ayacara stratigraphic sections; it indicates that most of the zircons are derived from igneous rocks (Rubatto 2002).
The majority of detrital zircons from the Ayacara
Formation’s upper member yield ages near 16 and
23 Ma (fig. 6) and show mostly a unimodal distribution in their relative probability age spectra. The
sample from the lower member similarly shows a
young age peak, but with the addition of much
older age peaks (Cretaceous ∼121–131 Ma, Carboniferous ∼303 Ma, and Proterozoic 1100–1200 Ma).
The maximum depositional ages for the full data
set range from ∼17.6 to 21.8 Ma (fig. 6).
Discussion
Age of the Ayacara Formation and Paleogeographic
Considerations. Planktic foraminifera studied here
indicate a Langhian (early Middle Miocene) age for
at least part of the upper member of the Ayacara
227
Formation (tables S1, S2). Accordingly, U-Pb geochronology shows that the youngest detrital zircon
populations from both the lower and the upper
members indicate an Early Miocene maximum age
(table S3; fig. 6). Considering that Ayacara Formation strata contain abundant volcanic material, it
is probable that sedimentation was contemporaneous with nearby volcanism, as previously suggested by Levi et al. (1966) and Rojas et al. (1994).
This implies that the younger zircon ages are likely
very close to the age of deposition. Thus, according
to the aforementioned data, the age of the Ayacara
Formation is Early-Middle Miocene. However,
there is no complete section of this unit, so it could
have a slightly wider age range. A minimum age
for the formation can probably be interpolated from
two radiometric ages (206Pb/238U in zircons) of
13.3 Ⳳ 0.2 Ma and 8.3 Ⳳ 0.4 Ma from dioritic porphyries and tonalite plutons at the Fiordo Pichicolo
and the Estuario Reloncavı́, respectively (figs. 1, 2),
which were correlated with dikes intruding the
Ayacara Formation (Duhart 2008). These ages must
be taken with some caution, as the analyzed samples were obtained from localities where crosscutting relationships between igneous and sedimentary rocks were not observed. Nevertheless, they
are in agreement with fission-track data (Adriasola
et al. 2006) and structural analysis (Cembrano et
al. 2000), indicating that significant uplift and exhumation commenced in this area during the Late
Miocene and likely put an end to deposition of the
Ayacara Formation. Thus, a minimum age of late
Middle-Late Miocene is inferred for the Ayacara
Formation.
The Early-Middle Miocene age for the Ayacara
Formation allows for some important paleogeographic inferences. According to these data, the
Ayacara Formation correlates with at least part of
the deep-marine Miocene strata of the Lacuı́ Formation exposed on Chiloé Island (see previous comment on the age of this unit; fig. 1) and equivalent
units that crop out along the forearc of south-central Chile (see Encinas et al. 2008, 2012b and references therein), and probably the Traiguén Formation to the south (Hervé et al. 2001). Subsurface
data from the Puerto Montt 1 well in the Longitudinal Valley (Elgueta et al. 2000) indicate that
below ∼2000 m of Plio-Pleistocene deposits there
is a Miocene marine succession that probably correlates with the Lacuı́ and Ayacara Formations. The
Miocene section overlies 1450 m of volcaniclastic
rocks that may also correlate with the Ayacara Formation (Katz 1965; Levi et al. 1966). The likely
source area, a contemporaneous volcanic arc, was
228
A. ENCINAS ET AL.
Figure 6. Relative probability age distribution plots of U-Pb ages (Ma) of detrital zircon from samples from the Caleta
Ayacara (CAYA) and Isla El Manzano (IMAN) sections.
likely close to the depositional basin, as suggested
by the abundance of conglomerate and coarsegrained sandstone in the Ayacara Formation. As
yet, no Miocene volcanics that could correlate with
the marine unit have been recognized in its vicinity. However, Early Miocene–Pliocene plutonic
rocks are abundant in the North Patagonian Andes
(fig. 1; Duhart 2008) because high exhumation rates
that commenced in the Late Miocene led to the
erosion of the Miocene volcanics and exposure of
the plutonic basement (see below).
U-Pb zircon geochronology can also give us important information regarding paleogeographic reconstructions of the study area during deposition
of the Ayacara Formation. Provenance inferences,
however, are hampered by the high exhumation
Journal of Geology
rates that affected the North Patagonian Andes
from the Late Miocene onward and gave way to the
erosion of the majority of the Meso-Cenozoic volcano-sedimentary cover rocks. Interestingly, there
are marked differences in detrital zircon ages between sample CAYA 4, from the lower shallowmarine basal member of the Ayacara Formation,
and samples IMAN 14, IMAN 20, CAYA 14, and
CAYA 24, all from the upper deep-marine member
(table S3, fig. 6). Sample CAYA 4 contains zircons
of different ages that include latest Oligocene–
Middle Miocene (major group between ∼17 and 24
Ma), Early Cretaceous (∼121–132 Ma), Early Jurassic (∼180 Ma), Carboniferous (∼303 Ma), Cambrian
(∼526 Ma), and Proterozoic (∼1100–1200 Ma) grain
populations. Carboniferous, Cambrian, and Proterozoic igneous rocks are unknown in the area, but
Duhart et al. (2009) found Carboniferous and Proterozoic detrital zircon populations in metamorphic rocks of the Patagonian Andes, which suggests
that they are recycled. Jurassic and Cretaceous plutonic and volcanic rocks occur only in the Main
Andean Cordillera, the latter solely exposed east of
the study area (Duhart 2008; fig. 1). Upper Oligocene–Middle Miocene igneous rocks are abundant
in the Main Andean Cordillera, although some
small outcrops of Upper Oligocene–Early Miocene
volcanic and subvolcanic rocks occur in the Coastal
Cordillera (fig. 1). Overall, these data suggest a
source area formed by rocks of different ages and
located in the Main Andean Cordillera, east of the
Ayacara Formation depositional basin.
In marked contrast, samples IMAN 14, IMAN 20,
CAYA 14, and CAYA 24, which are from two
different sections and belong to a stratigraphically younger interval, contain exclusively Late
Oligocene–Early Miocene zircon populations (table
S3, fig. 6). A possible explanation for this marked
change is that the Miocene transgression that
caused deposition of the deep-marine upper member of the Ayacara Formation completely covered
the former source area to the east and that sediment
provenance was limited to explosive volcanism of
a submerged volcanic arc (see Levi et al. 1966).
However, most of the conglomerate clasts in the
Ayacara Formation are subrounded, indicating
some degree of fluvial or alluvial transport (or wave
reworking) prior to their marine deposition. In addition, although most clasts are andesitic, they vary
in texture, which suggests that they did not come
from a single source but rather from the erosion of
different parts of the volcanic arc. If this volcanism
had been submarine, deposition of angular, compositionally and texturally similar clasts would be
expected. Other possibilities for the single Late
229
Oligocene–Early Miocene detrital zircon population is that the entire area was covered by the sea
except for an emerged volcanic arc or that an important episode of magmatism created large volcanic edifices that blocked river channels draining
Mesozoic-Paleozoic terrains. Another possible scenario is that a significant flux of Cenozoic zircons
simply diminished the older zircon signal. However, the high exhumation rates that affected the
study area during the Neogene have given way to
a very different geologic configuration between the
Miocene and the present day; thus, all of these possible explanations are speculative.
An important issue regarding paleogeographic
and tectonic reconstructions in the study area during the Miocene is the relationship between the
Ayacara Formation in the western Main Andean
Cordillera and the Cenozoic marine deposits of the
Rio Foyel and La Cascada Formations that crop out
only ∼70 km east in the eastern Andes (fig. 1).
Ramos (1982) related the Rio Foyel Formation and
equivalent deposits (including the La Cascada Formation) to a Pacific transgression and correlated
these strata with the Navidad-equivalent deposits
along Chile’s Pacific coast. However, the uncertainty over the age of these marine deposits has
complicated their correlation. The Rı́o Foyel Formation was ascribed to the Eocene on the basis of
molluscs (Chiesa and Camacho 2001), to the Early
Oligocene on the basis of 87Sr/86Sr (Griffin et al.
2004) and K/Ar dating (minimum age on crosscutting dikes by Giacosa et al. 2001), to the early Middle Oligocene–Early Miocene? on the basis of
foraminifers (Bertels 1980), and to the Late Oligocene–Early Miocene on the basis of palynomorphs
(Barreda et al. 2003) and foraminifers (Malumián et
al. 2008). The molluscan fauna of the La Cascada
Formation, on the other hand, has affinities with
the Eocene and Miocene faunas of Chile and Argentina according to Thiele et al. (1978), but their
poor preservation renders their identification at the
species level doubtful. Recently, Encinas et al.
(2012a) carried out LA-MC-ICPMS U-Pb geochronology on Rı́o Foyel and La Cascada detrital zircons
and obtained four maximum ages between ∼22 and
∼17 Ma. In addition, recent studies of calcareous
nannofossils from the Rı́o Foyel Formation indicated an Early-Middle Miocene age (Bechis et al.
2012). These results imply a temporal correlation
between the marine deposits from the eastern and
western flanks of the North Patagonian Andes.
However, it is unclear whether both successions
were deposited during the same marine transgression, as paleontologic evidence is contradictory. On
the basis of foraminifera and molluscs, Bertels
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A. ENCINAS ET AL.
(1980) and Griffin et al. (2002) favored an Atlantic
origin for the Rı́o Foyel Formation and a correlative
succession at Cerro Plataforma. However, Feldmann et al. (2011) and Ramos (1982) ascribed a Pacific origin for the Rı́o Foyel Formation on the basis
of its crustacean fossils and paleogeography. Finally, Nielsen et al. (2005) proposed an AtlanticPacific connection on the basis of his finding of a
gastropod of Atlantic origin, Perissodonta ameghinoi, in Miocene marine deposits near the Pacific
coast in the proximity of Valdivia (∼40⬚S). A significant problem is that most fossils from the eastern part of the Andean Cordillera are poorly preserved, which hinders their identification and
therefore makes it difficult to assess their origins.
Further study, including the recovery of better-preserved fossils, is needed to resolve this issue.
Possible Causes of Basin Subsidence. The sedimentologic, ichnologic, and micropaleontologic data
from this study allow reconstruction of the depositional environment in which the Ayacara Formation accumulated. The basal member is interpreted
as a fan-delta succession in which conglomeratic facies were deposited by episodic debris and hyperconcentrated flows and laminated sandstone reflects
wave reworking during “normal” marine sedimentation. The upper member is interpreted as a deepmarine succession where intermittent turbidity
flow deposits alternate with hemipelagic sedimentation. The presence of lower-middle bathyal benthic foraminifera in mixed-depth assemblages, as
well as the high planktic-to-benthic ratios, further
supports this interpretation (e.g., Van der Zwaan
1999). The occurrence of the Zoophycos ichnofacies
in fine-grained strata is also consistent with the micropaleontologic data (Frey and Pemberton 1984).
The abundance of gravel and sandstone versus mud
in these deposits, the predominance of basal Bouma
divisions, and the common occurrence of convolute
lamination in the turbiditic facies are characteristic
of tectonically active marine basins with narrow
shelves and high sedimentation rates (see Reading
and Richards 1994). Most of the sediment was likely
shed into the basin from a nearby volcanic arc, as
indicated by the predominance of volcanic clasts and
minerals and the abundance of interbedded tuffs.
Gravel and sand were probably transported short distances to the marine basin via alluvial fans, as indicated by the subrounded clasts. Deposition of waterlain tuffs would have resulted from ashfalls or
pyroclastic flows during or immediately following
explosive volcanism.
The Ayacara Formation represents a transgression culminating in deepwater deposition of its upper member. Because this member yielded some
species of benthic foraminifera that indicate lowermiddle bathyal depths (table S1) and maximum eustatic sea-level changes during the Cenozoic are on
the order of a few hundred meters (e.g., Haq et al.
1987), tectonic subsidence must have been the primary cause of this transgression. Forsythe and Nelson (1985) hypothesized that the Cenozoic Golfo
de Penas-Taitao basin, farther south near the Chile
Triple Junction (46⬚–48⬚S), could have formed as a
pull-apart basin in response to the northward displacement of the Chiloé block along the LOFZ.
This prompted some authors to speculate that the
Ayacara Formation and other supposedly correlative units between 42⬚S and 47⬚S were deposited in
a similar tectonic context (Pankhurst et al. 1992;
Hervé 1994; Rojas et al. 1994; Hervé et al. 1995),
but they provided little structural or chronological
evidence that relates dextral strike-slip motion on
the fault to coeval downwarping of these basins and
subsequent sedimentation therein (see Cembrano
et al. 2002). More importantly, correlation of the
Ayacara Formation with marine Miocene deposits
on Chiloé Island and in the Central Valley at 42⬚S
indicates that subsidence affected a wide area, ∼100
km from the present west coast of Chiloé to the
western Main Andean Cordillera. Furthermore,
Miocene subsidence not only affected the margin
at this latitude but almost the entire length of the
Chilean forearc (see Encinas et al. 2008). Therefore,
Miocene subsidence along the Chilean margin
must have been caused by a more regional event
than localized activity along the LOFZ or other tectonic discontinuities. Subsidence cannot be related
to the load imposed by a growing accretionary
wedge, as the actual prism in south-central Chile
is too small and probably formed within the last 2
m.yr. (Bangs and Cande 1997). In addition, subsidence in the study region affected an area many
tens of kilometers landward of the prism where the
upper plate is unaffected by accretionary loading.
Muñoz et al. (2000) and Jordan et al. (2001) described an episode of moderate extension that affected the Chilean margin between ∼33⬚S and 43⬚S
during the Late Oligocene–Early Miocene and resulted in extensional basins filled with thick successions of volcanic and sedimentary rocks from
the present Chilean coast to the retroarc. Muñoz
et al. (2000) ascribed extension to negative rollback
of the subducting slab caused by the transient
steepening of the subduction angle that they related
to a change from slower, more oblique South
America–Farallón convergence to more rapid, nearnormal South America–Nazca convergence at ∼25–
24 Ma (e.g., Pardo-Casas and Molnar 1987; Somoza
1998). Muñoz et al. (2000) also related forearc Mio-
Journal of Geology
cene marine sedimentation in south-central Chile
to this extensional event. However, their assumption is based principally on the idea that Miocene
marine rocks on northwestern Chiloé Island (fig. 1)
are interbedded with Late Oligocene–Early Miocene volcanic rocks of the Complejo Volcánico de
Ancud, but Garcı́a (1968) and Encinas et al. (2009)
demonstrated that these marine rocks unconformably overlie the volcanics and are therefore younger. In addition, most data indicate that extension
ceased in this region at ∼20 Ma and was followed
by contractional basin inversion (Muñoz et al.
2000; Jordan et al. 2001; Kay et al. 2006). Although
a small overlap exists, U-Pb and foraminiferal ages
obtained in this study indicate that the Ayacara
Formation commenced deposition coincident with
the end of the regional extensional event. Some
uncertainty remains, because a few radiometric
ages for some of the volcano-sedimentary successions believed to have been deposited under extensional tectonics are younger than 20 Ma (see Flynn
et al. 2008 and references therein).
On the basis of Von Huene and Scholl’s (1991)
study of convergent margins, Encinas et al. (2008,
2012b) related deep-marine Miocene sedimentation
and major forearc subsidence in south-central Chile
to subduction erosion, where the subducting plate
erodes the underside of the overriding plate, causing its thinning and margin subsidence. Melnick
and Echtler (2006) proposed that a low-relief Andes
resulted in a sediment-starved trench that, in addition to high plate-convergence rates, triggered
subduction erosion of the south-central Chilean
margin. Subsidence affected the entire width of the
forearc south of 38⬚ as well as the western flank of
the North Patagonian Andes at ∼42⬚S, in contrast
to other regions of Chile, where it was restricted
to the present coastal area. Yet, subduction erosion–driven subsidence is normally described for
offshore and coastal regions (Clift and Vannucchi
2004). South-central Chile’s anomalously thinned
crust, which was probably caused by the Oligocene–Early Miocene extensional tectonic event
(Jordan et al. 2001), could have facilitated tectonic
erosion–driven subsidence in this area (Encinas et
al. 2012b). Flattening of the slab, on the other hand,
has been proposed to cause basal erosion in forearc
areas from the trench to the depth of magma generation (see Kukowski et al. 2006 and references
therein). Muñoz et al. (2000) signaled that transient
steepening of the subduction angle could have
caused negative rollback in the south-central Chilean margin at ∼25 Ma. We speculate that subsequent slab flattening in combination with a sediment-starved trench, high convergence rates, and
231
an anomalously thinned crust could have caused
tectonic erosion and the related subsidence of the
entire forearc and western flank of the present
North Patagonian Andes.
On the basis of studies of the northeast Japan
margin and of numerical models of the interaction
between the subducting plate and the overriding
continent, Regalla et al. (2011) and Martinod et al.
(2012) independently proposed that increasing convergence rates can induce changes in slab geometry
and drive upper-plate subsidence. Considering that
a threefold increase in trench-normal convergence
rate between the Nazca and South American plates
took place at ∼25 Ma, this could be another cause
of forearc subsidence.
In summary, we consider subduction erosion to
be the most probable cause of forearc subsidence
in the North Patagonian Andes. We do not discard
as possible causes the extensional event described
by Muñoz et al. (2000) and Jordan et al. (2001) or
phenomena related to the interaction between the
oceanic and continental plates, such as changes in
the convergence rate. Additional research is needed
to better understand the vertical motion of forearcs
in convergent margins.
Implications for Andean Deformation. Ayacara
Formation marine strata deposited at bathyal
depths during the Early-Middle Miocene are tilted,
folded, faulted, and exposed in the western flank of
the North Patagonian Andes. Thus, after deposition
the unit uplifted, emerged, and contracted. Considering that the Ayacara Formation was deposited
in a deep-marine environment, Early-Middle Miocene is the maximum age for the deformation and
uplift of the western part of the Main Andean Cordillera in the study area. This implies that at least
the western part of the mountain chain formed
rather recently, which is a significant contribution
of this study. Our data agree with other studies of
the tectonic evolution of the Andean Cordillera in
the area ∼41⬚–42⬚S, although some uncertainties remain. The Early Miocene age (∼22–17 Ma) of the
Rı́o Foyel marine deposits and equivalent units
(Encinas et al. 2012a) is also the maximum age for
the uplift and tectonic deformation of the eastern
part of the Main Andean Cordillera. However, it is
not yet clear whether contraction and mountain
building were contemporaneous with marine deposition in this area. Orts et al. (2012) described
progressive unconformities in the lower Miocene
(∼18 Ma) marine succession of the Cerro Plataforma and in the probably coeval Late Oligocene?–
Middle Miocene? upper strata of the continental
Ñirihuau Formation, which they interpreted as evidence of ongoing contraction during sedimenta-
232
A. ENCINAS ET AL.
tion. This agrees with the ∼20 Ma age that most
authors consider the start of contractional basin
inversion in south-central Chile and Argentina
(Muñoz et al. 2000; Jordan et al. 2001; Kay et al.
2006). However, Bechis et al. (2011) cited growth
strata that they interpreted as synextensional in the
Troncoso Formation, a mostly continental sedimentary unit that overlies the Rı́o Foyel Formation
and was dated as 16.6 Ⳳ 0.5 Ma (U-Pb on zircons;
Bechis et al. 2012). Another uncertainty is the Pacific or Atlantic origin for the Rı́o Foyel Formation
and equivalent units (see above), which is difficult
to resolve because of the uncertain identifications
of its poorly preserved fossils and because its outcrops are absent to the west, along the central
North Patagonian Andes.
We envisage two possible scenarios for the tectono-sedimentary evolution of the Andean Cordillera in this region at ∼22–16 Ma. The first is that
a narrow and subdued mountain range rose as a
consequence of contractional inversion of basins
formed during a previous extensional phase (Muñoz
et al. 2000; Jordan et al. 2001; Kay et al. 2006). This
incipient cordillera separated two distinct Pacific
and Atlantic marine transgressions or presented
narrow seaways that permitted either a limited
connection between the two seas or the incursion
of the Pacific to the east of this range. In the second
scenario, the previously described Oligocene–Early
Miocene extensional tectonic phase continued in
this region with marine deposition. This implies
that the Andean Cordillera had not yet formed,
which likely facilitated either the transoceanic connection or the incursion of the Pacific in the Rı́o
Foyel area. An emerging volcanic arc would have
formed a series of islands between the eastern and
western flanks of the present cordillera. According
to most published data, a marine connection between the Atlantic and the Pacific in the first tectonic scenario is more likely.
The angular unconformity between the Ñirihuau
Formation and the overlying Collón-Curá Formation clearly evidences tectonic Andean uplift and
contractional tectonics in the eastern flank of the
North Patagonian Andes during the Middle to Late
Miocene (Bechis and Cristallini 2005). Carbon and
oxygen isotope studies indicate significant uplift of
the southern Patagonian Andes at ∼16.5 Ma (Blisniuk et al. 2005). During the Late Miocene, deformation probably ceased in the eastern Main Andean Cordillera and was displaced to the western
part (Folguera and Ramos 2002), where it has been
related to the dextral strike-slip LOFZ (Cembrano
et al. 2002; Thomson 2002; Adriasola et al. 2006).
Evidence for this important tectonic discontinuity
comes from structural analysis and 40Ar/39Ar geochronology from mylonites at 41⬚–46⬚S, which
show a Late Miocene–Pliocene dextral strike-slip
event, oblique-slip, and contractional deformation
indicative of bulk transpressional deformation
(Cembrano et al. 2000, 2002). These data agree with
zircon and apatite fission-track thermochronology
that supports denudation of the southern Andes at
the same latitudes during the Late Miocene–
Pliocene, related to dextral transpressional activity
in the LOFZ combined with intense erosion by
Plio-Pleistocene glaciation (Thomson 2002; Adriasola et al. 2006). Enhanced exhumation gave way
to erosion of the majority of the Meso-Cenozoic
volcano-sedimentary cover and widespread exposure of the plutonics that form the North Patagonian Batholith (Adriasola et al. 2006). Cembrano et
al. (2002) and Thomson (2002) proposed a crustalscale transpressional pop-up structure defining the
main range of the Patagonian Cordillera, principally based on the spatial distribution of geologic
units, fission-track data that show deeper crustal
levels in the central part of the Main Andean Cordillera, more abundant Meso-Cenozoic cover rocks
in the borders of this orogen, and the oblique-slip
and contractional shear zones in the eastern and
western parts of the main Patagonian Batholith.
This pop-up structure consists of several blocks
characterized by different amounts of cooling and
denudation and separated by major oblique- or
reverse-slip faults that link at depth below the main
trace of the Liquiñe-Ofqui fault (Thomson 2002).
Subsurface geologic data from the submerged Longitudinal Depression west of the Main Andean Cordillera agrees with this model. The ENAP Puerto
Montt 1 well records Miocene marine strata covered by ∼2000 m of Plio-Pleistocene deposits
(Elgueta et al. 2000). The 4010-m well did not reach
basement even though it was only ∼20 km west of
the western Main Andean Cordillera, where deformed Miocene marine strata of the Ayacara Formation crop out and Meso-Cenozoic plutonic rocks
are widely exposed. Additional information from
ENAP seismograms from the Golfo de Ancud, between Chiloé Island and the area where the Ayacara
Formation crops out, shows subhorizontal strata,
presumably Cenozoic, in the western Longitudinal
Depression that become folded to the east (González 1983). Thus, in agreement with the Cembrano
et al. (2002) and Thomson (2002) pop-up model, we
infer a significant tectonic discontinuity between
the Longitudinal Depression and the Main Andean
Cordillera that separates an uplifted, strongly deformed domain to the east from a downward,
mildly deformed domain to the west (fig. 7). Flex-
Journal of Geology
233
Figure 7. Cross sections showing the interpreted tectono-sedimentary evolution of the study area during the Neogene.
See the explanation in the text at the end of “Discussion.” The part of the figure corresponding to the Andean
Cordillera has been modified from Thomson (2002).
ural subsidence induced by the major uplift of the
Main Andean Cordillera, in combination with significant glacial erosion, gave way to 2000 m of PlioPleistocene synorogenic deposits in the eastern part
of the Longitudinal Depression (fig. 7).
With regard to the tectonic activity of the LOFZ
in the study area, the age of this major discontinuity
is uncertain, and it is not yet clear whether the activity of this fault zone is restricted to the Late Miocene to present that most age data infer (e.g., Cembrano et al. 2000, 2002; Thomson 2002; Adriasola et
al. 2006) or whether it began prior to the Late Cre-
234
A. ENCINAS ET AL.
taceous and was reactivated at various times in response to an evolving plate-kinematic framework
(Cembrano et al. 2000). The cause of the transpressional tectonics associated with the LOFZ has also
been debated. Some studies speculated that oblique
subduction was the driving mechanism (Hervé 1977;
Beck 1988). However, Forsythe and Nelson (1985)
considered the angle of oblique convergence in the
region (∼25⬚) to probably be small to solely account
for the strike-slip activity of the LOFZ, and they
alternatively proposed the indenter effect of the
Chile Ridge at the southern termination of the fault
system. Cembrano et al. (2000) deduced that oblique
subduction combined with strong intraplate coupling were the key factors in the long-term motion
along the LOFZ. According to them, ridge collision
is a second-order mechanism that has favored dextral displacement along the southern portion of the
LOFZ only since the Pliocene. Whereas plate motions and convergence angle have not changed substantially during most of the Cenozoic, partitioning
of the Nazca–South America slip vector into forearc
shortening and intra-arc dextral-strike displacement
may have been continuous for most of this period
(Cembrano et al. 2000). On the basis of the lack of
apparent denudation between 42⬚S and 46⬚S before
∼16 Ma that was indicated by fission-track data,
Thomson (2002) discarded transpressional tectonics
in this area before the Middle Miocene. Because the
Chile Ridge first collided with the Pacific margin of
the South American plate at ∼14 Ma and very rapid
cooling and denudation rates in the southern Andes
between ∼7 and 2 Ma are concurrent with the collision of several short segments of this spreading
center, Thomson (2002) concluded that the indenting of the Chile Ridge acted as a major driving force
in sustaining the late Cenozoic dextral transpressional tectonism. Similar conclusions were reached
farther north (41⬚–42⬚S) by Adriasola et al. (2006),
also on the basis of fission-track data.
Our data indicate that the onset of major regional
subsidence and deep-marine sedimentation in the
Ayacara area (i.e., very close to the main trace of
the LOFZ) occurred during the Early-Middle Miocene, which implies that transpression in the area
started afterward, in accordance with conclusions
based on fission-track data (Thomson 2002; Adriasola et al. 2006). Therefore, motion along the LOFZ
is probably linked to the arrival of the Chile Ridge
and either did not commence before ∼14 Ma or was
dominated earlier by horizontal or transtensional
strike-slip tectonism, as suggested by Thomson
(2002).
Finally, integrating our data with published information enables us to propose a model (sum-
marized in fig. 7) for the Neogene tectono-sedimentary evolution of the study area. First, regional
subsidence, probably related to subduction erosion,
affected the entire forearc and at least the western
part of the Main Andean Cordillera during the
Early-Middle Miocene. Subsidence initiated a significant transgression from the west, and deepwater
deposition ensued. An incipient growing Cordillera
probably blocked the Atlantic transgression that
reached the eastern flank of the Andes in the Rı́o
Foyel area from reaching the Pacific ingression, although narrow seaways probably connected the
two oceans. Second, during the Late Miocene tectonic activity along the LOFZ led to the initial development of a transpressional pop-up structure
that uplifted the Main Andean Cordillera and resulted in deformation and emergence of the Miocene marine Ayacara Formation strata. It is not
clear whether subsidence and deep-marine sedimentation was still active in the forearc to the
west, as the age of the marine deposits in that area
is under revision (see above). Third, transpressional
deformation, uplift, and exhumation in the Main
Andean Cordillera have continued from the Pliocene to the present day. Uplift and emergence of
the Neogene marine successions in the Coastal
Cordillera and western Longitudinal Depression
also occurred during this interval. Their uplift has
been attributed to underplating of sediments because their strata show little deformation and no
evidence of significant tectonic shortening (Encinas
et al. 2008). In the eastern part of the Longitudinal
Depression, on the other hand, flexural subsidence,
ascribed to the uplift of the Main Andean Cordillera
and its intense denudation, gave way to deposition of ∼2000 m of Plio-Pleistocene synorogenic
deposits.
Conclusions
First, marine deposits in the western flank of the
North Patagonian Andes at ∼42⬚S, termed the Ayacara Formation, were deposited in a deep-marine
environment during the Early-Middle Miocene.
Second, correlation of the Ayacara Formation with
coeval deep-marine deposits in the present Coastal
Cordillera and Longitudinal Depression indicate
that marine deposition was caused by a major regional event of tectonic subsidence probably related to subduction erosion that affected the entire
forearc and the western flank of the Andean Cordillera. Third, an Early-Middle Miocene age for the
Ayacara Formation is a reliable maximum for the
onset of deformation and uplift of the western part
of the North Patagonian Andes, in agreement with
Journal of Geology
previous studies based on microstructural analysis
and fission-track thermocronology. Fourth, major
regional subsidence and deep-marine sedimentation of the Ayacara Formation close to the main
trace of the LOFZ during the Early-Middle Miocene implies that transpressional tectonics associated with this fault system commenced after the
Middle Miocene, in accordance with fission-track
data reported by others.
235
ACKNOWLEDGMENTS
This research was funded by Fondecyt (Fondo Nacional de Desarrollo Cientı́fico y Tecnológico) projects 11080115, 3060051, and 1110914 of Conicyt
(Comisión Nacional de Investigación Cientı́fica y
Tecnológica) and the Institut de Recherche pour le
Développment. L. A. Buatois also received financial
support from Natural Sciences and Engineering Research Council Discovery Grants 311726-05/08.
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