Major forearc subsidence and deep-marine Miocene sedimentation

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Major forearc subsidence and deep-marine Miocene
sedimentation in the present Coastal Cordillera and Longitudinal
Depression of south-central Chile (38°30′S–41°45′S)
Alfonso Encinas1,†, Kenneth L. Finger2, Luis A. Buatois3, and Dawn E. Peterson2,*
1
Departamento de Ciencias de la Tierra, Universidad de Concepción, Casilla 160-C, Concepción, Chile
University of California Museum of Paleontology, Valley Life Sciences Building 1101, Berkeley, California 94720, USA
3
Department of Geological Sciences, University of Saskatchewan, 114 Science Place, Saskatoon S7N 5E2, Canada
2
ABSTRACT
INTRODUCTION
Neogene marine strata crop out in the
present Coastal Cordillera and Longitudinal Depression of south-central Chile between 38°30′S and 41°45′S, indicating the
onset of a major marine transgression that
covered most of the forearc in this area. In
order to determine the sedimentary environment, paleobathymetry, and age of these
deposits, we carried out integrated sedimentologic, ichnologic, and micropaleontologic studies on samples from oil wells
and outcrops in the region. Our results indicate that these successions were deposited
at lower-bathyal depths (>2000 m) during
the middle to late Miocene. Shallow-marine
deposition followed in the southwestern part
of the study area during the Pliocene(?). We
attribute deep-marine sedimentation in this
area to a major event of subsidence in the
Miocene that affected the entire forearc and
that was caused by basal subduction erosion. We suggest that the anomalously thin
crust that characterizes this area may have
facilitated forearc subsidence and allowed
the Miocene transgression to advance much
farther inland here than in other regions of
Chile. Subsequent uplift of the forearc is
ascribed to basal accretion or underplating
of sediments. Our conclusions contradict
previous studies that favor a stable margin
at these latitudes since the Jurassic. Deepmarine sedimentation in this area during the Miocene implies that the present
Coastal Cordillera and Longitudinal Depression were probably submerged during
that epoch.
Forearc regions are among the most active
sites related to plate tectonics, being frequently
associated with destructive megathrust earthquakes and subjected to rapid and pronounced
uplift or subsidence (e.g., Plafker and Savage,
1970; von Huene and Scholl, 1991; Rehak et al.,
2008). Evidence for large-scale vertical motion
in forearc areas comes principally from offshore
and onshore sedimentologic and micropaleontologic studies that reveal the onset of significant
events of alternating subsidence and uplift (e.g.,
von Huene and Scholl, 1991; Buret et al., 1997;
Chanier and Ferriére, 1999; Melnick and Echtler,
2006; Finger et al., 2007; Encinas et al., 2008b).
Convergent margins are classified as accretive
or erosive based on their material-transfer modes
(e.g., von Huene and Scholl, 1991; Clift and Vannucchi, 2004). Accretive margins form in regions
where trench sediment thickness exceeds 1 km or
where there is a slow convergence rate, and they
are characterized by net frontal growth due to
accretion and forearc uplift caused by the underplating of sediment (Clift and Vannucchi, 2004).
Erosive margins form where trench sediment
thickness is less than 1 km and there is a slow
convergence rate, and they are characterized by
trench retreat, extensional faulting, and regional
forearc subsidence driven by tectonic erosion
(von Huene and Scholl, 1991; Clift and Vannucchi, 2004). Thus, trench fill and plate-convergent
rate exert first-order controls on forearc tectonics. Other factors that may also play important
roles in the type of convergent margin that forms
include inherited upper-plate structures, strain
partitioning, and subduction of lower-plate
topographic highs. Additional studies, however,
are needed to further our understanding of the
controls and dynamics of deformation in forearc
regions (Rehak et al., 2008).
†
E-mail: [email protected]
*Deceased
The Chilean margin, which is more than
4000 km long, has been an area of ongoing subduction of oceanic crust at least since the Jurassic (Mpodozis and Ramos, 1989), and there are
historic records of megathrust earthquakes, such
as the 1960 Valdivia (or Great Chilean) earthquake, which was the largest ever recorded,
registering 9.5 on the Richter scale (e.g., Plafker
and Savage, 1970). Thus, this convergent margin is one of the most important natural laboratories in the world in which to understand
forearc dynamics, both present and past. Two
different segments can be distinguished in this
margin based on their material-transfer modes.
The north Chilean margin (~18°S–34°S) is a
classic example of a sediment-starved margin (von Huene and Ranero, 2003) that has
undergone subduction erosion during the last
~200 m.y., as deduced by the continuous migration of the magmatic arc since the Jurassic (Rutland, 1971). In contrast, the margin of
south-central Chile (38°S–45°S) shows a welldeveloped accretionary wedge, and apparently
has never been subjected to subduction erosion,
since no major systematic changes in the location of the magmatic belt have occurred in this
region since the Late Jurassic (Pankhurst et al.,
1999). Yet, Bangs and Cande (1997) argued
that the small dimensions of the present wedge
could indicate the onset of alternating episodes
of accretion and tectonic erosion in the past (see
following). The presence of Upper Cretaceous,
Eocene, and Neogene marine sequences in the
forearc of south-central Chile may be evidence
of such alternating episodes in which subduction erosion has forced subsidence and marine
transgressions, whereas regressions could reflect intervals dominated by accretion and uplift.
Among these marine sequences, the most
widespread is that deposited during the Neogene
along almost the entire length of the Chilean
GSA Bulletin; July/August 2012; v. 124; no. 7/8; p. 1262–1277; doi: 10.1130/B30567.1; 8 figures; Data Repository item 2012128.
1262
For permission to copy, contact [email protected]
© 2012 Geological Society of America
Major forearc subsidence and deep-marine Miocene sedimentation, south-central Chile
forearc, from Iquique (~20°13′S) to Golfo the
Penas (~47°30′S; see Encinas et al., 2008b,
and references therein). The reference unit of
these deposits is the Navidad Formation, which
is well exposed in the coastal bluffs of southcentral Chile (~34°S; Encinas et al., 2006). Recent sedimentologic and paleontologic studies
(Finger et al., 2007; Encinas et al., 2008b, and
references therein) have confirmed that the Miocene strata exposed along the coast of southcentral Chile are deep-marine deposits. The
Miocene beds also crops out inland, particularly
in the region between ~38°S and 41°S (Figs. 1
and 2), where the sea transgressed much farther
eastward and covered an area that extends from
the present coastline nearly to the western limit
of the Andean range. In order to understand the
causes of this important extension, the age and
sedimentary environment of these deposits must
be accurately determined. However, there are
significant discrepancies between the conclusions of previous sedimentologic and micropaleontologic studies on these strata. Based
on benthic foraminifers, Martínez-Pardo and
Zúñiga (1976) inferred outer-neritic to upperbathyal depths of 100–300 m for Neogene
deposits in the northern part of the area, near
Temuco (Figs. 1 and 2), whereas Osorio and
Elgueta (1990) concluded that deposition was
at lower-bathyal depths below 2000 m. In the
southern part of the area, between Valdivia and
Puerto Montt (Figs. 1 and 2), sedimentologic
analyses by Elgueta et al. (2000) and le Roux
and Elgueta (2000) suggested that deposition
took place at much shallower depths in coastal
embayments. In contrast, foraminiferal analyses
by Martínez-Pardo and Pino (1979), Marchant
and Pineda (1988), and Marchant (1990) indicated outer-shelf to upper-bathyal depths,
SOUTH
AMERICA
TEMUCO
39°S
N
Paci
fi
40°S
c Oc
ean
0 10 20 30 40 50 km
Pleistocene-Holocene marine
and continental deposits
Pliocene? marine rocks of the
Caleta Godoy Formation
Middle to late Miocene marine rocks
of the Santo Domingo Formation
VALDIVIA
Cenozoic volcanic rocks of the
Main Andean Cordillera
CC
Early Oligocene to early Miocene volcanic
rocks of the “Coastal volcanic belt”
ID
MAC
Eocene? coal-bearing rocks
OSORNO
Paleozoic and Mesozoic
plutonic rocks
Paleozoic metamorphic rocks
PUERTO
MONTT
74°W
73°W
LOFZ
41°S
72°W
Figure 1. Geologic map of the study area, modified from Elgueta (1990), Sernageomin (1998), Antinao et al. (2000),
le Roux and Elgueta (2000), Duhart et al. (2003), and Rehak (2008). CC—Coastal Cordillera; ID—Intermediate
Depression; MAC—Main Andean Cordillera ; LOFZ—Liquiñe-Ofqui fault zone.
Geological Society of America Bulletin, July/August 2012
1263
Encinas et al.
A
Victoria
SOUTH
AMERICA
0
10
B
20 km
A
Cholchol
TEMUCO
Cunco
39°S
50 km
Nueva Imperial
Cu1
0
10
20
30 km
Paci
fi
PUERTO
MONTT
74°W
Maullín
72°W
20 km
CT1 Ct-B1
La Unión
Cg2
20 km
RCH1
Rb1
Tg1
D
Cg1 Lago
Lm1
Llanquihue
O
RN
SO
O
Hm2
10
10
Rh1
Be1
Hm1
D
0
0
Rh2
Bahía
Mansa BM1
BM2
41°S
Santo Domingo
Section
Lago Collico
C
OSORNO
PC1
PC2
Punta Hueicolla
Osorno
Section
C
CSD1
CSD2
PH1
B
VALDIVIA
VALDIVIA
Corral
Cunco
c Oc
ean
25
TEMUCO
Lb5 Lb6
N
0
CR1
LP1
Lb2,4,7,8,10
Ch2
Ch1
Playa Colún
Puerto
Saavedra
Estaquillas
PO1 P. Ortiga
Section
PQ1
Punta
Quillagua
Lm2
PM1
PUERTO
MONTT
Maullín
Carelmapu
Figure 2. (Left) Map of study area showing the location of outcrops of the Santo Domingo Formation (light gray) and the Caleta Godoy
Formation (dark gray). (Center and Right, A–D) Detailed maps showing locations of ENAP boreholes (open circles), sampled outcrops
(solid squares), and sections (solid triangles) cited in the text, Tables DR1–DR6 (see text footnote 1), and Figure 2. Well name abbreviations:
LP—Los Pinos; Ch—Cholchol; Lb—Labranza; Ct-B1—Catamutun-B1; Rh—Rahue; Be—Bellavista; Hm—Huilma; Rb—Rio Blanco;
Tg—Tegualda; Cg—Colegual; Lm—Los Muermos; PM—Puerto Montt. Labranza 2, 4, 7, 8 and 10 are not differentiated due to their proximity. Mapping of Santo Domingo and Caleta Godoy Formations is modified from Elgueta (1990), Sernageomin (1998), Antinao et al. (2000),
and Duhart et al. (2003).
whereas Cecioni (1970) referred to the middlebathyal depths (1000–2000 m) interpreted by
Natland and Severin from ENAP (The Chilean
National Petroleum Company) well samples.
Discerning the sedimentary environment and
paleobathymetry of these deposits is vital to
understanding the paleogeography and tectonics of the south-central Chilean forearc during
the Neogene. There are several key questions to
address here: (1) Was the cause of marine transgression in this area eustatic or tectonic, and
if the latter, what events triggered subsidence
and subsequent uplift of the margin? (2) Was
the forearc of south-central Chile completely
submerged during any part of the Neogene?
(3) Why did the Neogene sea almost reach the
present Main Andean Cordillera in this area but
not to the north, where correlative deposits are
1264
known only from the coastal area? We attempt
to answer these questions here by integrating
stratigraphic, sedimentologic, ichnologic, and
micropaleontologic data. The main objectives of
this contribution are to discern the sedimentary
environment and age of these Neogene deposits and to explore their tectonic and paleogeographic implications.
GEOLOGIC SETTING
The study area is located in the forearc
of south-central Chile, between the cities of
Temuco (38°44′S) and Puerto Montt (41°28′S)
(Figs. 1 and 2). At these latitudes, the oceanic
Nazca plate and the continental South American
plate converge at a rate of 62 mm yr –1 (Kendrick
et al., 2003). Three physiographic units trend-
ing N-S are distinguished in this area; from
west to east, these are the Coastal Cordillera,
the Longitudinal Depression, and the Main
Andean Cordillera. The Coastal Cordillera is a
subdued mountain range that parallels the Chilean trench. The Longitudinal Depression, also
referred to as the Intermediate Depression or the
Central Valley, is a low-lying area situated between the Coastal Cordillera and the Main Andean Cordillera. The Main Andean Cordillera is
distinguished by having the highest peaks in the
region and active volcanoes. In the study area,
the Coastal Cordillera is largely formed by a
basement made up of Paleozoic metamorphic
rocks and minor Paleozoic and Mesozoic plutonic rocks (Sernageomin, 1998) (Fig. 1). Tertiary sedimentary and volcanic strata, as well
as Pleistocene and Holocene unconsolidated
marine and continental deposits, locally cover
basement rocks (Sernageomin, 1998) (Fig. 1).
Generally thick and extensive vegetation and
soil cover hinder geologic mapping in the Longitudinal Depression. However, information
obtained from outcrops in road cuts and river
banks, as well as from ENAP and mining boreholes, indicates the presence of units similar to
those of the Coastal Cordillera (Sernageomin,
1998; Elgueta et al., 2000). Cenozoic sedimentary and volcanic rocks predominate in most of
the Longitudinal Depression, although hill belts
formed by Paleozoic metamorphic rocks transect the Central Valley and join the Coastal and
Main Andean Cordilleras in some areas (Fig. 1).
The paucity of good outcrops has resulted in
the proposal of several stratigraphic schemes for
the Tertiary succession exposed in the Coastal
Cordillera and the Intermediate Depression
(e.g., Brüggen, 1950; García, 1968; Sernageomin, 1998). Cisternas and Frutos (1994) simplified the Tertiary stratigraphy by defining three
basic units—a lower continental coal-bearing
unit, a middle volcanic unit, and an upper marine
fossiliferous unit. Although the ages of some of
these units are debatable (see following), wells
drilled in the Temuco area confirm their relative
stratigraphic position (Rubio, 1990).
The basal coal-bearing unit consists of conglomerate, breccia, sandstone, siltstone, and
coal (Sernageomin, 1998). This unit overlies the
Paleozoic metamorphic basement and was deposited in alluvial and estuarine environments
according to le Roux and Elgueta (2000). The
age of these strata had been considered as Oligocene or Miocene (e.g., Sernageomin, 1998), but
a recent study by Finger and Encinas (2009) reassigned them to the Eocene based on the recognition of planktic foraminifers belonging to the
genus Globigerinatheka in the absence of any
younger species.
Volcanic rocks of the middle unit sporadically crop out in the study area. They are part of
the Mid-Tertiary Coastal volcanic belt formed
during a widespread volcanic episode that radiometric dating places in the late-early Oligocene
to early Miocene interval (Muñoz et al., 2000).
These rocks overlie the coal-bearing unit in the
Temuco area (Rubio, 1990).
The marine fossiliferous section that is the
upper unit and main focus of this study was
originally named as the Navidad Formation in
the pioneering work of Brüggen (1950). Subsequent workers redefined these deposits with several local names, notably the Chochol Formation
(García, 1968) in the Temuco area, the Santo
Domingo Formation (Martínez-Pardo and Pino,
1979) and the Huilma Formation (Céspedes and
Johnson, 1984) in the Valdivia-Osorno area,
and the Lacui Formation (Valenzuela, 1982) in
the area located around Puerto Montt and north
of Chiloé Island. Our studies reveal geographically separate successions of similar sedimentary
facies and microfossils; therefore, we now refer
to all of them as the Santo Domingo Formation,
which is the better known and most extensive
of the units (e.g., Sernageomin, 1998; Elgueta
et al., 2000). Outcrops of this formation occur
in the coastal area, Coastal Cordillera (mostly
its western and eastern flanks), and Longitudinal Valley. In the Temuco area, outcrops of
the Santo Domingo Formation occur almost to the
eastern edge of the Central Valley and near
the Main Andean Cordillera (Figs. 1 and 2). This
unit unconformably overlies the Paleozoic metamorphic basement (Sernageomin, 1998), the
coal-bearing deposits of the lower unit (Illies,
1970), and the volcanic deposits of the middle
unit (Rubio, 1990). The Santo Domingo Formation has a maximum thickness of 1500 m based
on oil wells (Elgueta et al., 2000), although the
stratigraphic thickness of any individual outcrop
is no more than 100 m. This unit has a basal
transgressive interval of breccia, conglomerate,
or sandstone overlain by sandy siltstone, sandstone, and minor breccia (García, 1968; Illies,
1970). Planktic foraminifers suggest a middle to
late Miocene age for this unit (see following for
details and discussion).
Although Cisternas and Frutos (1994) referred to the Miocene Santo Domingo Formation as the upper unit, we include a fourth
marine unit that overlies the latter. This succession was defined as the Caleta Godoy Formation by Valenzuela (1982) and crops out in
the southwestern part of the study area (Figs. 1
and 2). Nevertheless, it is possible that it also
extends to other areas, such as Valdivia, where
some outcrops show facies characteristic of this
unit. However, the lack of fossils and the bad
preservation of these outcrops preclude their
accurate stratigraphic identification. The Caleta
Godoy Formation overlies the metamorphic
Paleozoic basement (Valenzuela, 1982) and the
Santo Domingo Formation. It consists of sandstone and minor conglomerate and contains fossil molluscs (Antinao et al., 2000). Age-dating
studies have yet to be performed on this unit, but
its stratigraphic position suggests a Pliocene age
(see following for discussion).
In the northern limit of our study area
(~38°S), the Andean Cordillera exhibits a major
geomorphologic change as the broad and high
(>4000 m) Central Andes to the north give way
to the narrow and relatively low (~2000 m)
North Patagonian Andes to the south (Hervé,
1994). In addition, the Andes south of 38°S
have a crustal thickness of only ~40 km (e.g.,
Lüth et al., 2003), which amounts to only 15 km
of tectonic shortening, and they lack an active
foreland fold-and-thrust belt (Melnick et al.,
2006). Compressional deformation in the North
Patagonian Andes only occurred during the Late
Cretaceous, late Eocene, and late Miocene, alternating with intervals of tectonic relaxation
(Zapata and Folguera, 2005). The oldest rocks
of the Main Andean Cordillera are middle–late
Paleozoic metamorphic rocks (Duhart, 2008).
Late Jurassic–Miocene plutonic rocks of the
North Patagonian Batholith are the dominant
rocks in the southern part of the Andean range
at the studied latitudes (Mpodozis and Ramos,
1989). Paleogene and Neogene volcanic, volcaniclastic, and sedimentary rocks become
more abundant in the northern part of this range
(Glodny et al., 2008). A significant tectonic discontinuity in the Main Andean Cordillera east of
the study area is the N-S–trending, dextral strikeslip, intra-arc Liquiñe-Ofqui fault zone, which
extends from the Chile triple junction at 46°S to
the Copahue volcano at 38°S (e.g., Hervé, 1994).
The Liquiñe-Ofqui fault zone decouples the
south-central Chilean forearc, which behaves
as a northward-gliding sliver (Cembrano et al.,
2002; Melnick et al., 2009). High-strain zones
in plutonic rocks and metamorphic wall rocks
within the Liquiñe-Ofqui fault zone resulted
from dextral strike-slip deformation accompanied by syntectonic pluton emplacement during
the Miocene–Pliocene (Cembrano et al., 2002,
and references therein). The Liquiñe-Ofqui fault
zone appears to be linked to a zone of magmatic
weakening of the crust, as suggested by alignment of the strike with Pliocene to Holocene
volcanic centers (Glodny et al., 2008).
MICROPALEONTOLOGIC ANALYSIS
In order to determine the age and paleobathymetry of the Santo Domingo Formation,
we studied foraminifera and ostracodes. We collected samples from 12 outcrops, and borrowed
slides from 23 ENAP wells drilled in the study
area (Fig. 2; Tables DR1−DR61).
Biostratigraphy
The vast majority of planktic and benthic
foraminifers identified in the present study
(Tables DR1–DR4 [see footnote 1]) also occur in the late Miocene–early Pliocene fauna
of the Navidad, Ranquil, and Lacui Formations
that crop out in the coastal areas of Navidad
(~34°S), Arauco (~37°S), and Chiloé (~42°S),
respectively (Finger et al., 2007). However, a
1
GSA Data Repository item 2012128, Micropaleontologic Data, is available at http://www.geosociety
.org/pubs/ft2012.htm or by request to editing@
geosociety.org.
1265
Encinas et al.
refined age span for the Santo Domingo Formation remains elusive, as the unit yields few
planktic foraminifers, and index species are
scarce. All of the planktic foraminifers recovered in this study from strata of the Santo
Domingo Formation (Tables DR1–DR2 [see
footnote 1]) occur in the Miocene (Tables DR5
and DR6 [see footnote 1]). Some of the species
ranges extend older or younger than Miocene,
but three localities yielded assemblages with
concurrent ranges that refine the age to a subepoch: (1) the Labranza 7 well has Globorotalia
(Fohsella) fohsi fohsi Cushman and Ellisor, restricted to middle Miocene planktic foraminifer
zone N12; (2) the Tegualda 1 well has two species that first appear in the late Miocene (zone
N16), Globorotalia (Globoconella) conomiozea
Kennett and Neogloboquadrina pachyderma
(Ehrenberg); and (3) the PQ outcrop yielded
Globoturborotalita woodi–Gt. apertura transitional forms that suggest a late Miocene age
(Kennett and Srinivasan, 1983). The occurrence in the Tegualda 1 well of the robust late
Eocene−early Miocene (zones P13−N6) species
Catapsydrax dissimilis Cushman and Bermúdez
is attributed to reworking. Thus, the foraminiferal fauna of the Santo Domingo Formation appears to be middle to late Miocene. However,
we cannot discard a broader age range because
of the paucity of planktic markers. Samples
from outcrops of correlative successions in the
areas of Navidad, Arauco, and Chiloe, as well
as from the ENAP H well drilled into the continental shelf off Valdivia, yielded many more
specimens, including rare and relatively minute
markers for the early Pliocene such as Globorotalia (Globoconella) puncticulata (Deshayes)
(Finger et al., 2007; Encinas et al., 2008b).
A revision of former foraminiferal studies
carried out in the study area reveals similar results to those attained in our study (see Tables
DR2, DR4, and DR6 [see footnote 1]). In the
Temuco area, Osorio and Elgueta (1990) studied the Labranza 1 and Cunco 1 boreholes and
reported planktic foraminifers indicative of
a middle Miocene to early late Miocene age.
However, their faunal list includes Globigerina
druryi Akers, a late Miocene species (zones
N16−N17) and several specimens with their
first occurrences in the Tortonian (zone N16).
If their identifications are accurate, the early
to middle Miocene planktic foraminifers also
recorded in this association must have been reworked in the late Miocene. Farther south, in
the area around Valdivia and Osorno, MartínezPardo and Pino (1979), Marchant and Pineda
(1988), and Marchant (1990) indicated a middle
Miocene age for the Santo Domingo Formation,
in part based on benthic foraminifers. This is a
problematic approach because bottom-dwelling
1266
organisms are primarily facies controlled, and
therefore their ranges tend to be interbasinally
time-transgressive (Ingle, 1980). MartínezPardo and Pino (1979) recorded eight species of planktic foraminifers near Valdivia, but
none was firmly identified, as all were modified with “cf.” (confer with). Among the comparable species they listed, Globigerinoides
trilobus (Reuss) and Globorotalia continuosa
(Blow) have a concurrent range zone of N9−
N16 (middle to late Miocene). In the Catamutun
area, Marchant and Pineda (1988) and Marchant (1990) reported Globigerina pachyderma,
which first occurs in N16 (Tortonian, late Miocene) and suggests that the Santo Domingo Formation is late Miocene in age.
There are no paleontologic data that reveal
the age of the Caleta Godoy Formation. Stratigraphic superposition dictates that it is not older
than late Miocene age. However, in the areas of
Navidad (34°S) and Arauco (37°S), shallowmarine units ascribed to the Pliocene overlie the
deep-marine late Miocene−early Pliocene Navidad and Ranquil Formations (Martínez-Pardo
and Osorio, 1968; Encinas et al., 2006). It is
therefore likely that the Caleta Godoy Formation correlates with the cited nearshore units that
are of Pliocene age.
Paleoenvironmental Indications
Benthic foraminifers are the most useful
means of determining paleobathymetry because
of their great abundance and diversity in modern seas, and the availability of their recorded
depths in many regions. Because most benthic displacement is downslope, upper-depth
limits of foraminiferal species are considered
minimum depths of deposition (Ingle, 1980).
Whereas the major water masses control both
vertical and geographic distributions of offshore
benthic foraminifers, greater accuracy is likely
to be attained if the fossil and modern faunas are
in the same region.
The modern upper-depth limits used in this
study were derived from the data recorded by
Bandy and Rodolfo (1964), Ingle et al. (1980),
and Resig (1981) on foraminiferal species currently living along the Pacific margin of central
South America. The upper-depth limits of extant
or morphologically similar congeneric species
were extrapolated to the fossil assemblages, and
the deepest upper-depth limit represented was
taken as a minimum depth of deposition. Where
mass wasting has occurred, the deepest-dwelling species might be relatively rare components of a mostly displaced assemblage. Thus,
shallow-water species commonly dominate assemblages recovered from the middle and lower
continental slope (Ingle, 1980).
The paleodepth methodology used in this
study is not infallible, however, because of its
inherent assumption that the vertical distribution
of water masses in the region has not changed
significantly from the Miocene to the present.
Nevertheless, the employment of upper-depth
limits (minimum depths) is conservative, and
the unknown inaccuracies of the depth determinations are unlikely to be significant in the overall reconstruction of depositional history.
All of the studied assemblages are mixeddepth associations of shelf and slope species
(Tables DR3–DR6 [see footnote 1]); thus,
downslope displacement is evident. The deepestdwelling taxa in the majority of samples indicate minimum depths of deposition on the lower
slope (2000–4000 m). Only three assemblages
have their deepest dwellers with shallower
upper-depth limits, all in the upper-middle
bathyal (500−1500 m) zone (Tables DR3–DR4
[see footnote 1]); two of those assemblages,
however, are relatively weak (i.e., consisting of
few specimens). Species indicative of lowerbathyal depths are Bathysiphon spp., Cornuspiroides foliacea (Philippi), Ellipsopolymorphina
schlichti (Silvestri), and Melonis pompilioides
(Fichtel and Moll) f. sphaeroides Voloshinova,
and species indicative of lower-middle bathyal
depths are Osangularia bengalensis (Schwager),
Pyramidulina acuminata (Hantken), Siphonodosaria hispidula (Cushman), and S. sagrinensis
(Bagg), several of which are similar or identical
to those van Morkhoven et al. (1986) included in
the cosmopolitan deep-water fauna.
The vast majority of ostracode species found
in this study are endemic to the Miocene of the
southeast Pacific, although some also occur in
the Caribbean and southwest Atlantic. Six taxa
are referred to the psychrospheric fauna (Tables
DR3–DR4 [see footnote 1]), which Benson
(1975) described as characteristic of the deep
(≥500 m) cold-water masses. Their association
with shallower-dwelling ostracodes supports the
foraminiferal evidence in revealing that most of
these are mixed-depth associations.
STRATIGRAPHIC, SEDIMENTOLOGIC,
AND ICHNOLOGIC FINDINGS
In the field, we performed detailed sedimentologic and ichnologic analyses on outcrops of
the Santo Domingo and Caleta Godoy Formations. Exposures in the study area are small and
scarce, limited to just a few coastal cliffs and
road cuts. This is due to the prevailing climate,
which promotes a thick vegetative cover and
weathering and erosion of the soft sedimentary rocks. As a consequence, we could only
measure partial sections up to ~100 m in thickness (Fig. 3). Where possible, we measured
Height
in section
(m)
CUESTA SANTO DOMINGO
Facies
PUNTA ORTIGA
80
?
50
OSORNO
(PO1)
SD2
Facies
10
SD3
(CSD2)
SD1?
20
SANTO DOMINGO FORMATION
30
SD1
40
?
SD2
60
70
CG2 CG4 CG2
CALETA GODOY FORMATION
Facies
0
GS
abcdefghijk
LEGEND
Lithology
Sedimentary structures
Fossils
Trace fossils
Conglomerate
Trough cross-bedding
Foraminifers
Chondrites
Breccia
Planar cross-bedding
Ostracodes
Zoophycos
Parallel lamination
Bivalves
Phycoshiphon
Gastropods
Ophiomorpha
Crustaceans
Thalassinoides
Sandstone
Siltstone
Type of contact
Angular unconformity
Echinoderms
Wood
Leaves
Planolites
Asterosoma
Terebellina
Figure 3. Representative sections of the study area from Punta Ortiga, Osorno, and Santo Domingo (see Fig. 2 for location
of sections). Grain size (GS): a—clay, b—silt, c—very fine sandstone, d—fine sandstone, e—medium sandstone, f—coarse
sandstone, g—very coarse sandstone, h—gravel (granules), i—gravel (pebbles), j—gravel (cobbles), k—gravel (boulders).
Facies are indicated in the logs. Location of microfossil samples PO1 and CSD2 (see Tables DR2, DR4, and DR6 [see text
footnote 1]) in Punta Ortiga and Santo Domingo sections is also indicated. The thickness of the Santo Domingo Formation
at Punta Ortiga is probably underestimated because this unit is strongly deformed in this area. The Santo Domingo section
and the lower half of the Osorno section are partially covered by vegetation.
1267
Encinas et al.
bedding attitudes: strata of the Santo Domingo
Formation generally dip 0°–30° in different
directions, while those of the Caleta Godoy
Formation are always horizontal. A particularly
interesting succession is that of Punta Ortiga
(Figs. 1–4), where Santo Domingo strata are
strongly faulted and inclined up to 90°, resulting in an angular unconformity with the overlying Caleta Godoy Formation.
Santo Domingo Formation (SD)
Exposures of this unit occur along the entire
study area. Because of the outcrop limitations,
it was impossible to measure a complete section of this unit. The basal or upper contact of
this formation can only be observed in a few
locations.
SD Facies Descriptions
SD facies 1—basal conglomerate and sandstone. A basal conglomerate/breccia occurs at
Bahía Mansa, Corral, and Punta Ortiga (Fig. 2).
It overlies metamorphic basement in the latter two localities, but its basal contact is not
exposed at Punta Ortiga (Fig. 3). This facies
is represented by up to a few-tens-of-meters
thick, clast-supported, poorly to moderately
sorted conglomerate and breccia with subrounded to angular schistose clasts that range
from 10 cm to 1 m in size. It locally contains
molluscan shell fragments. The basal conglomerate is overlain by a fining-upward succession
of interbedded conglomerate/breccia and sandstone that transitions into sandy-siltstone of SD
facies 2 at Bahía Mansa and Corral. At those
localities, the conglomerate is sharp to erosive
basally, clast to matrix supported, and poorly to
moderately sorted. It contains schist and quartzite clasts, millimeters to decimeters in size, and
typically in random orientation. Sandstone
is medium to very coarse grained and locally
shows cross-bedding and parallel lamination.
Bioturbation is represented by Thalassinoides.
Conglomerate and sandstone commonly show
erosive contacts.
At Punta Ortiga, the succession is a few tensof-meters thick and consists of a basal conglomerate overlain by well-sorted medium-grained
sandstone that grades upward into sandy siltstone of SD facies 2 (Fig. 3). Sandstone shows
a massive aspect that is probably due to pervasive faulting having obliterated any sedimentary
structures. However, thinly interbedded sandstone and siltstone occur in the basal interval,
which contains body fossils of bivalves, gastropods, and decapods. Beds are locally bioturbated, with BI = 1–3 (bioturbation index of
Taylor and Goldring, 1993). Asterosoma and
Phycoshiphon dominate the ichnofauna, with
Thalassinoides, Ophiomorpha, and Chondrites
as accessory elements (Figs. 3 and 5A). Wood
logs with Teredolites are locally present. At
Osorno (Fig. 2), the basal part of the section
consists of sandstone and conglomerate (Fig. 3)
that may correspond to SD facies 1, but this is
uncertain due to insufficient exposure.
SD facies 2—sandy siltstone. This is the most
typical facies of the Santo Domingo Formation.
A
B
Figure 4. (A) Contact (indicated by arrow) between the Santo Domingo Formation and the overlying Caleta Godoy Formation at Punta Ortiga
(see Figs. 2 and 3). Strata of the Santo Domingo Formation are inclined (not evident in this photograph but visible in B), whereas strata
of the Caleta Godoy Formation are horizontal. The photo shows an ~25 m section of the cliff face. (B) Basal breccia of the Santo Domingo
Formation (right) grading upward into well-sorted sandstone (left) at Punta Ortiga. Encircled rucksack for scale is ~50 cm high. Angular
clasts that constitute the breccia can be observed in a fallen block lying on the beach in the lower part of the photo. Strata are vertical to
slightly overturned (indicated by white lines). This point is located ~100 m from that of image A.
1268
B
A
D
C
Figure 5. Characteristic trace fossils of the Santo Domingo Formation (see Figs. 2 and 3). (A) Asterosoma isp. in basal shallowmarine sandstones of the Santo Domingo Formation at Punta
Ortiga. (B) Chondrites isp. at Hueicolla. Coin for scale is 2 cm in
diameter. (C) Phycosiphon isp. at Punta Ortiga. (D) Zoophycos
isp. at Punta Ortiga. Scale bar = 1 cm. B, C, and D are typical
ichnofacies of deep-marine sandy siltstones of the same unit.
Continuous successions of sandy siltstone can
be at least 25 m thick. This facies is dark gray
when fresh, and slightly fissile, probably due
to small differences in grain size. It is mainly
composed of quartz and muscovite. No clear
primary features were observed, likely due to
intense bioturbation (see following). It locally
shows scarce intercalations of fine- to mediumgrained sandstone. The fossil biota is diverse
and consists of bivalves, gastropods, brachio-
pods, bryozoans, crustaceans, echinoids, fishes,
foraminifers, ostracodes, radiolarians, calcareous nannoplankton, and leaves (Mai, 1982;
Chirino-Gálvez, 1985; Pino and Beltrán, 1979;
Martínez-Pardo and Pino, 1979; Covacevich
et al., 1992; Kutscher et al., 2003; Nielsen,
2005). Intense bioturbation (BI = 5–6) locally
shows abundant Chondrites, Zoophycos, and
Phycosiphon, and very scarce “Terebellina”
(Figs. 5B–5D). At Cuesta Santo Domingo (near
Valdivia), the succession includes a monospecific association of Chondrites (Encinas et al.,
2008a) as well as other trace fossils resembling
Phycosiphon and Zoophycos.
SD facies 3—rhythmically interbedded sandstone and siltstone. Medium- to coarse-grained
sandstone and minor pebbly conglomerate are
rhythmically interbedded with siltstone in this
facies (Fig. 6A). Bed thickness ranges from
centimeters to decimeters and rarely to meters.
1269
Encinas et al.
A
B
C
Figure 6. Characteristic facies of the Santo Domingo Formation (see Figs. 2 and 3). (A) Rhythmically interbedded
sandstone and siltstone near Osorno. Scale (white line) is
2 m. (B) Parallel-laminated sandstone that grades upward
into an interval with convolute lamination at Playa Colun.
Coin for scale is 2.5 cm in diameter. (C) Sheared flame structure in sandstone at Playa Colun.
Most sandstones have sharp, or less commonly
erosive, contacts. Sedimentary structures are
not always evident due to weathering, but parallel lamination and convolute bedding, locally
showing an upward transition between both
structures (Fig. 6B), can be observed. Load and
flame structures are also present at the contact
of some strata (Fig. 6C). Sandstone is composed
of quartz and muscovite with minor plagioclase,
amphibole, and biotite. Siltstone contains foraminifers and Chondrites.
SD facies 4—conglomerate. This facies consists of decimeter- to meter-thick conglomerate/
breccia strata interbedded with sandy siltstone
bearing lower-bathyal benthic foraminifers (SD
facies 2) in Bahía Mansa. Conglomerate is sharp
to erosive basally, clast to matrix supported, and
1270
moderately to poorly sorted. Clasts are quartz
and schist, subrounded to angular, centimeters
to decimeters in size, and usually in random
orientation.
SD Facies Interpretation
SD facies 1. The basal conglomerate is a
nearshore facies deposited during the initial
stage of the Miocene transgression. Interbedded
conglomerate and sandstone at Bahía Mansa
and Corral are interpreted as fan-delta successions. Conglomerate was deposited by episodic
debris flows, as indicated by their poor to moderate sorting, randomly orientated clasts, and
abundant matrix (e.g., Nemec and Steel, 1984;
Hwang and Chough, 1990). The presence of
Thalassinoides in the associated cross-bedded
sandstone clearly indicates marine influence.
The most likely depositional scenario is uppershoreface multidirectional littoral currents reworking the sediment supplied by the fan-delta
system (Clifton, 2006; Komar, 1976). The sandstone succession overlying the basal breccia at
Punta Ortiga is interpreted as recording deposition in a transgressive shoreface. The trace-fossil
association and the vertical facies relationships
suggest a middle- to lower-shoreface setting
(Buatois and Mángano, 2011). The abundance
of wood fragments with Teredolites also points
to a transgressive context (Savrda et al., 1993).
The presence of rhythmically interbedded laminae of sandstone and siltstone in some parts of
this succession may suggest tidal influence (Nio
and Yang, 1991).
SD facies 2. Dark-gray siltstones can be
deposited in a wide variety of low-energy settings. However, the presence of mixed-depth
associations of benthic foraminifers and ostracodes in this facies evidences downslope displacement and final deposition of sediments
on the lower continental slope (see previous
section). Thus, this facies was likely deposited
by distal turbidity currents and hemipelagic settling. Interbedded sandstones are interpreted as
coarser-grained turbidites. The trace-fossil association is ascribed to the Zoophycos ichnofacies,
which occurs in quiet-water settings with low
oxygen levels associated with high concentrations of organic detritus, and it is typical of
the outer shelf and slope (Frey and Pemberton,
1984). The dark-gray color of this facies and
the occurrence of pyrite (Chirino-Gálvez, 1985)
also indicate a substrate rich in organic matter
and low in oxygen. Thick, fine-grained intervals within deep-marine successions are usually interpreted as condensed sections deposited
at low sedimentation rates during the time of
maximum regional transgression (Weimer and
Slatt, 2006). The abundant and diverse biota and
the high content of organic matter present in SD
facies 2 are also characteristic features of condensed sections (Weimer and Slatt, 2006).
SD facies 3. Although the relatively poor
preservation of sedimentary structures precludes a more refined interpretation, the rhythmic nature of this facies and the presence of
sandstones with parallel lamination and convolute bedding, which sometimes are transitional,
are characteristic of turbidites with Bouma divisions (Bouma, 1962; Posamentier and Walker,
2006). Load-and-flame structures observed at
the base of some sandstone beds are also typical of turbidites and indicate rapid deposition
(e.g., Posamentier and Walker, 2006). However,
none of these sedimentary structures is exclusive to turbidites, and rhythmically interbedded
sandstone and siltstone can form in other environments (e.g., tidally influenced settings and
fluvial overbanks). Yet, the presence of mixeddepth associations of shelf and slope benthic
foraminifers in this facies evidences downslope
displacement and final deposition at lowerbathyal depths (see previous section), which
strongly support our interpretation. The presence of monospecific ichnofossil assemblages
of Chondrites suggests oxygen-depleted conditions and deposition in relatively deep water
(Bromley and Ekdale, 1984).
SD facies 4. We attribute this facies to debris flows based on its moderate to poor sorting,
random orientations of clasts, and abundant
matrix, which together suggest that cohesive
strength was the dominant mode of transport
(e.g., Nemec and Steel, 1984; Hwang and
Chough, 1990). The presence of interbedded
sandy siltstone with lower-bathyal benthic foraminifers indicates deep-marine deposition of
this facies.
Caleta Godoy Formation (CG)
Exposures of the Caleta Godoy Formation
are limited to the southwestern part of the study
area, where good outcrops only occur in coastal
bluffs. Sedimentologic and ichnologic analyses of the unit’s exposed sections were limited
because only their basal part and fallen blocks
were accessible. Nonetheless, sedimentary
structures and trace fossils are exceptionally
well preserved in the accessible rocks. This unit
overlies the Santo Domingo Formation and the
metamorphic Paleozoic basement. The contact
between the Caleta Godoy and Santo Domingo
formations is rarely observed, and, depending
on the locality, it is conformable or unconformable. The angular unconformity seen at Punta
Ortiga (Figs. 3 and 4) is evidence that the underlying Santo Domingo Formation had been uplifted, deformed, and partially eroded prior to
subsidence and deposition of the Caleta Godoy
Formation. Some outcrops of the Caleta Godoy
Formation reveal a thin basal conglomerate
characterized by well-rounded, centimeter- to
decimeter-sized clasts of schist and quartzite,
and locally include siltstone clasts from the underlying Santo Domingo Formation.
CG Facies Descriptions
CG facies 1—sandstone with planar lamination. This facies consists of medium- to
coarse-grained sandstone with low-angle planar
or parallel lamination, which locally caps crossbedded sandstone (CG facies 2).
CG facies 2—cross-bedded sandstone and
conglomerate. Medium- to very coarse-grained
sandstone, pebbly sandstone, and conglomerate constitute this facies (Figs. 7A–7B), which
exhibits trough and planar cross-bedding. Conglomerate is clast-supported and well sorted,
containing subrounded to well-rounded, centimeter- to decimeter-sized clasts of schist and
quartzite, and is commonly interbedded with
coarse-grained sandstone. Bioturbation is absent or limited to low-density Skolithos burrows.
Wood fragments and coal laminae are locally
present.
CG facies 3—lam-scram sandstone. Decimeter-thick interbedded strata of structureless
and hummocky cross-stratified fine- and very
fine-grained sandstone (colloquially called
“lam-scram” after Howard, 1975) characterize
this facies (Fig. 7C). Structureless (scrambled)
sandstone is highly bioturbated and shows
abundant Ophiomorpha and Planolites and
minor Schaubcylindrichnus, Phycosiphon, and
“Terebellina” (Fig. 7D). Hummocky intervals show sets with low-angle foresets dipping
in different directions and basal erosive contacts with onlapping laminae. Deep Ophiomorpha burrows are the only trace fossils present in
these intervals.
CG facies 4—fine-grained bioturbated sandstone. The deposits are structureless, highly
bioturbated sandstone with Ophiomorpha and
Planolites (Fig. 7E).
CG Facies Interpretation
CG facies 1. The presence of planar-laminated sandstone in vertical association with
upper-shoreface, cross-bedded sandstone of CG
facies 2 indicates deposition in the intertidal
foreshore area of a wave-dominated shoreline
and reflects the high energy of intense swash and
backwash (e.g., Clifton, 2006). The absence of
bioturbation is consistent with the high-energy
conditions of the foreshore (MacEachern and
Pemberton, 1992; Buatois and Mángano, 2011).
CG facies 2. This facies records deposition from multidirectional current flows typical of the buildup and surf zones of the upper
shoreface (Clifton, 2006). Its association with
well-sorted conglomerate having well-rounded
clasts also indicates continuous water agitation.
The presence of a monospecific low-density
association of Skolithos sp. is consistent with
continuous migration of large bed forms, commonly resulting in sparse colonization by the
benthic fauna. This association represents
the Skolithos ichnofacies, which is typical of,
but not exclusive to, shallow-marine environments of relatively moderate to high energy
(MacEachern and Pemberton, 1992; Buatois
and Mángano, 2011).
CG facies 3. The lam-scram facies results
from the alternation of storm and fair-weather
processes, where storm waves cut an erosional
surface, which is then overlain by low-angle,
laminated-to-hummocky, cross-stratified sand
(Harms et al., 1975). Following the abatement
of storm conditions, an opportunistic suspension-feeding fauna colonizes the upper part
of the sandy unit, forming dwelling structures
(Pemberton et al., 2001; Mángano et al., 2005).
During fair weather, when energy is low, intense bioturbation from the activity of climax
populations obliterates the primary sedimentary fabric. Lam-scram sandstone occurs in the
lower shoreface of wave-dominated shorelines
(MacEachern and Pemberton, 1992; Buatois
and Mángano, 2011). In particular, lam-scram
deposits are typical of moderately stormaffected shorefaces of intermediate energy
(MacEachern and Pemberton, 1992; Mángano
et al., 2005; Buatois et al., 2007).
1271
Encinas et al.
A
B
C
D
E
Figure 7. Characteristic facies and trace fossils of the Caleta Godoy
Formation (see Figs. 2 and 3). (A) Interbedded cross-bedded sandstone and conglomerate at Punta Ortiga. (B) Trough and planar
cross-bedded sandstone at Punta Quillagua. (C) Lam-scram sandstone at Carelmapu. (D) Highly bioturbated structureless (scrambled) sandstone with abundant Ophiomorpha isp. (note pelletal
wall of specimen just above the tip of the pen) and minor Planolites
isp. at Carelmapu. (E) Ophiomorpha isp. and Planolites isp. in finegrained bioturbated sandstone at Estaquillas.
CG facies 4. The intense bioturbation in this
facies indicates relatively slow sedimentation
rates in a low-energy setting (Buatois and Mángano, 2011). However, the fine-grained sandy
texture and the association with nearshore facies
suggest deposition above the fair-weather wave
1272
base. Integration of sedimentologic, ichnologic,
and stratigraphic evidence suggests deposition
in a lower-shoreface environment during fair
weather. In particular, CG facies 4 corresponds
to a weakly storm-affected lower shoreface of
low energy (MacEachern and Pemberton, 1992).
DISCUSSION
Sedimentologic, ichnologic, and micropaleontologic data obtained in this study permit reconstruction of the geologic history of
the forearc of south-central Chile during the
Neogene (Fig. 8). A middle to late Miocene
marine transgression covered most if not all
of the study area from the present coast to the
eastern part of the Longitudinal Valley, and deposited strata of the Santo Domingo Formation.
A flooding surface/sequence boundary (FS/SB)
or co-planar surface formed at the base of this
unit due to amalgamation of erosion during
lowstand and the subsequent transgression.
A lag of coarse-grained deposits produced by
ravinement associated with the sea advance
mantled this surface. Afterward, a fining- and
deepening-upward succession formed. Sediments characteristic of fan-delta and open-coast
clast environments were deposited, forming the
basal part of the Santo Domingo Formation.
The upper interval of this unit was deposited
at bathyal depths where deposition was by
episodic turbidity currents and the continual
settling out of suspension of fine hemipelagic
particles. Subsequently, strata of the Santo
Domingo Formation were gently folded, uplifted, emerged, and eroded. A new transgressive cycle commenced in the Pliocene(?) with
deposition of the Caleta Godoy Formation. An
FS/SB surface formed at the base of this unit
and was covered by a transgressive conglomerate. Sedimentologic and ichnologic features
in the Caleta Godoy Formation indicate deposition in a wave-dominated, shallow-marine clastic environment. The presence of large wood
fragments and coal intercalations suggests a
wave-dominated deltaic system. In accordance,
the relatively simple ichnofacies that characterizes lower-shoreface facies of this unit could be
due to periodic fluvial discharge into the marine environment. Marine deposition in the region terminated with the uplift of the Neogene
section, which rose hundreds of meters in some
parts of the present Coastal Cordillera and Longitudinal Depression.
Considering that benthic foraminifers of the
middle to late Miocene Santo Domingo Formation indicate depositional water depths below 2000 m for part of this unit, and maximum
1) MIDDLE-LATE MIOCENE
Andean
Cordillera
Forearc
Sea level
PACIFIC PLATE
Subduction erosion
2) PLIOCENE (?)
Andean
Cordillera
Forearc
Sea level
PACIFIC PLATE
Figure 8. Cross sections showing the interpreted tectonosedimentary evolution of the
Chilean forearc in the study area during
the Neogene. (1) During the middle to late
Miocene, a major event of subsidence related to subduction erosion gave way to an
important marine transgression that covered most of the forearc and deep-water
deposition of the Santo Domingo Formation. (2) During the Pliocene(?), the forearc
was uplifted by underplating of sediments,
prior extensional basins were inverted, and
strata of the Santo Domingo Formation
were gently folded, emerged, and partially
eroded. Subsequently, a new transgressive
cycle gave way to shallow-marine deposition
of the Caleta Godoy Formation. (3) After
the Pliocene, forearc uplift continued but
was concentrated in the coastal area, where
it created the Coastal Cordillera. Strata of
the Santo Domingo and the Caleta Godoy
Formations were uplifted and emerged,
being completely eroded in the highest summits of this range.
3) PRESENT DAY
Underplating
Coastal
Cordillera
Longitudinal
Depression
Andean
Cordillera
Sea level
PACIFIC PLATE
Underplating
LEGEND
Basement rocks
Caleta Godoy
Formation
Subducted sediments or rocks
removed by subduction erosion
Santo Domingo
Formation
Quaternary
deposits
Frontally and basally accreted
sediments
1273
Encinas et al.
global sea-level changes during the Neogene
are just in the order of a few hundred meters
(Haq et al., 1987), the Miocene transgression
in this region of south-central Chile must have
been predominantly forced by regional tectonics. The catalyst would have been a major event
of tectonic subsidence that affected the entire
width of the forearc, from the present coast to
the eastern Longitudinal Valley. As noted already, deep-marine Miocene deposits also occur
north and south of this region (Encinas et al.,
2008b). The major event of tectonic subsidence
therefore must have affected the whole length of
the Chilean margin. We dismiss causes for subsidence that imply widespread extension, such as
an increment in the rollback of the subduction
hinge relative to the advance of the overriding
plate, because such events would affect not only
the continental margin but also the Main Andean
Cordillera. On the contrary, it was a compressive
tectonic phase coeval with deep-marine sedimentation in the forearc that affected the Andes
of south-central Chile during the middle–late
Miocene (see following). We also discard the
explanation of transtensional tectonics caused
by dextral strike-slip displacement along the
Liquiñe-Ofqui fault zone because deep-marine
Miocene strata in this region extend across an
~100-km-wide area (i.e., some deposits are far
from the area affected by this fault zone). In
addition, correlative successions that crop out in
other areas of Chile occur beyond the longitudinal extent of this fault zone. Thus, the cause
of the regional Miocene subsidence that affected
the entire Chilean margin cannot be accounted
for by local tectonic discontinuities. Based on
previous studies of convergent margins (e.g., von
Huene and Scholl, 1991), Encinas et al. (2008b)
proposed basal tectonic or subduction erosion as
the most likely mechanism. By this process, removal of the underside of the upper plate results
in its thinning and subsidence of the margin (von
Huene and Scholl, 1991). As indicated by the
presence of deep-marine Miocene deposits near
the western flank of the Andes, subsidence affected the entire width of the forearc in the study
area, in contrast to other regions of Chile, where
it was restricted to the present coastal area. We
suggest that the difference could be due to the
significant reduction in crustal thickness south of
38°S (Lüth et al., 2003) that was caused by an
extensional tectonic event during the Oligocene−
early Miocene (Jordan et al., 2001; see following). Tectonic erosion of an anomalously thinned
crust could have facilitated forearc subsidence in
the study area.
Subduction erosion therefore appears to be
the most probable cause of margin subsidence
in south-central Chile during the Miocene. Yet,
data based on temporal arc migration for this
1274
area contradict those based on the subsidence
record. Assuming a constant slab angle, the
eastward shift of volcanic arcs in Chile has
been attributed to subduction erosion (e.g.,
Rutland, 1971). However, although there is
evidence of eastward arc migration during the
Neogene in northern (Rutland, 1971), central
(Kay et al., 2005), and southernmost Chile
(Cande and Leslie, 1986), no major systematic changes in the location of the magmatic
belt have taken place between 39°S and 47°S
since the Late Jurassic (Pankhurst et al., 1999).
Bangs and Cande (1997), on the other hand,
argued that the small dimension of the present
accretionary wedge at these latitudes suggests
the onset of alternating episodes of accretion
and tectonic erosion in the past. On the basis of
mass balance, they indicated that the size of the
actual prism is inconsistent with a long history
of accretion and probably formed in the last
1–2 m.y. A similar line of reasoning was followed by Glodny et al. (2006), who indicated
that the margin of south-central Chile must have
shifted between accretionary and tectonic-erosional modes within short intervals not allowing for major net growth or mass loss through
time. Margin subsidence data presented in this
study could be evidence of one of the erosional
phases speculated by these authors.
Nearshore strata of the Pliocene(?) Caleta
Godoy Formation overlying deep-marine deposits of the Miocene Santo Domingo Formation implies that the margin was significantly
uplifted between the deposition of these units.
As previously noted, the angular unconformity
observed between these units in some parts
of the study area indicates that the Santo Domingo Formation had to have been deformed,
emergent, and partially eroded prior to deposition of the Caleta Godoy Formation. Continued
uplift of the forearc led to the definitive emersion of the Santo Domingo and Caleta Godoy
Formations, which were elevated up to some
hundreds of meters above sea level as observed
in some parts of the present Coastal Cordillera
and the Longitudinal Valley. Yet, recent uplift
of the forearc did not reach similar magnitudes
along the entire study area, as the northern
(~37°S−38°30′S) and southern (~40°S−41°S)
segments of this region rose during the Quaternary while the central segment (~38°30′S−
40°S) was in a quasi-stable state (Rehak et al.,
2008). Late Cenozoic uplift of Neogene marine
successions in the Navidad (34°S) and Arauco
(37°S) areas has been attributed to underplating of sediments (Melnick and Echtler, 2006;
Encinas et al., 2008b). After studying ENAP
seismic lines from the areas offshore of Arauco
and Valdivia, Melnick and Echtler (2006) proposed tectonic inversion of Miocene extensional
basins and syntectonic deposition of nearshore
Pliocene deposits. Previous studies of seismic
images in our study area also show tectonic
inversion of originally normal faults (Elgueta
et al., 2000; Jordan et al., 2001). In accordance
with the geophysical data, field work carried out
in this study observed the angular unconformity
between the horizontal beds of the Pliocene(?)
Caleta Godoy Formation and the tilted strata of
the Miocene Santo Domingo Formation.
Encinas (2006) and Melnick and Echtler
(2006) speculated that Miocene tectonic erosion
and subsidence followed by Pliocene underplating and uplift of the margin along the Navidad
(34°S) and Arauco (37°S) areas could have been
caused by a change in trench sediment flux induced principally by climatic changes. Currently, thick trench deposits along the humid
southern Chilean margin correlate with accretion, whereas thin trench sediment in the arid
northern Chile margin correlates with nonaccretion or tectonic erosion (Bangs and Cande,
1997). This agrees with the observations of
Clift and Vannucchi (2004), who studied different convergent settings and noted that accretive
margins form in regions where trench sediment
thickness exceeds 1 km, whereas erosive margins characterize areas with a thinner trench fill.
Trench sediment flux depends principally on
coastal precipitation rates, as these determine the
amount of continental erosion (Lamb and Davis,
2003). Trench sedimentation rates are significantly augmented during glacial periods through
enhanced denudation of elevated areas (von
Huene and Scholl, 1991). Based on these observations, Encinas (2006) speculated that a shift
from a more arid to a more humid climate could
have caused a transition from tectonic erosion
to sediment underplating along the margin of
central Chile during the Pliocene. Melnick and
Echtler (2006) proposed that a low-relief Andes
resulted in a sediment-starved trench, which, in
addition to high plate-convergence rates, caused
subduction erosion of the south-central Chilean
margin during the Miocene. The subsequent uplift and the onset of glaciation in the Patagonian
Andes during the latest Miocene–early Pliocene
(Mercer and Sutter, 1982; Rabassa and Clapperton, 1990; Rabassa et al., 2005) resulted in
denudation and increased sediment flux to the
trench, causing basal accretion and forearc uplift
from the Pliocene onward (Melnick and Echtler,
2006). The model proposed by these authors is
intriguing and could account for the subsidence
and subsequent uplift of the forearc in south-central Chile. However, Miocene–Pliocene major
subsidence and subsequent uplift of the forearc
also took place north of 33°S (see Encinas et al.,
2008b, and references therein), where no significant glaciation occurred during that time, and,
as noted by Clift and Hartley (2007), the present
trench is largely devoid of sediment. The latter
authors speculated an alternative cause—that
the inversion could be related to changes in plate
coupling caused by slowing convergence rates.
The uncertainties highlight the necessity of further work to fully understand the vertical motion
of convergent margins.
Data presented in this study, integrated with
published information on the evolution of the
North Patagonian Andes, are important in order
to understand the geologic history of this region
during the Neogene. Most authors agree that an
important extensional event, attributed to negative rollback by Muñoz et al. (2000), affected
the area between ~33°S and 43°S during the late
Paleogene–early Neogene (Jordan et al., 2001;
Burns et al., 2006). Extensional basins filled by
thick successions of volcanic, fluvial, and lacustrine rocks extended across a broad area from
the present Chilean coast to the retroarc (Muñoz
et al., 2000; Jordan et al., 2001). In the forearc
of south-central Chile, these mostly volcanic
and volcaniclastic deposits are referred to as the
Mid-Tertiary Coastal magmatic belt (Muñoz
et al., 2000). In the Andean Range, volcanic and
sedimentary deposits accumulated in the CuraMallín basin between 36°S and 39°S (Jordan
et al., 2001), and in the Ñirihuau basin between
40°S and 42°S (Spalletti and Dalla Salda, 1996).
In some cases, K/Ar and Ar/Ar radiometric
ages obtained for these successions in different
subbasins of south-central Chile and Argentina do not overlap, which could be due either
to diachronous sedimentation or, more likely,
to limitations of some of the radiometric dates
as previously discussed by Flynn et al. (2008).
However, most data place deposition of these
strata during the late Oligocene–early Miocene,
and extending into the middle Miocene in some
areas (see Muñoz et al., 2000; Jordan et al., 2001;
Flynn et al., 2008). After this extensional period,
the Main Andean Cordillera and the forearc of
south-central Chile experienced a different tectonic evolution during the middle–late Miocene,
when a compressive phase affecting the Andean
range resulted in shortening, uplift, exhumation,
and inversion of the former basins (Melnick
et al., 2006; Folguera et al., 2010). In contrast,
major subsidence of the forearc resulted in deepmarine deposition, as indicated by the strata
of the Santo Domingo Formation superjacent
to volcanic rocks of the Mid-Tertiary Coastal
magmatic belt (Rubio, 1990). Subsequently, the
Main Andean Cordillera experienced a different
evolution in the northern and southern parts of
the study area. Whereas transpressional deformation has been concentrated exclusively along
the dextral strike-slip Liquiñe-Ofqui fault zone
at ~41°S–42°S from the late Miocene to pres-
ent (Folguera and Ramos, 2002; Adriasola et al.,
2006), extension has dominated at ~39°S since
the early Pliocene (Folguera et al., 2010). The
forearc of south-central Chile, on the other hand,
has been rising since the Pliocene, as discussed
herein. This is in accordance with previous studies by Hartley et al. (2000) on the Neogene geologic history of northern Chile, which also found
the tectonostratigraphic evolution of the forearc
to differ from that of the Andean range.
Our findings also have important implications for understanding the paleogeographic
and geomorphologic evolution of the forearc
of south-central Chile during the late Cenozoic.
Exposures of the Santo Domingo Formation
occur in the region extending from the coast
to near the western limit of the Main Andean
Cordillera. The lower-bathyal (>2000 m) depositional depth indicated by benthic foraminifers
for this unit implies that the present Coastal
Cordillera and Longitudinal Valley were likely
submerged during the middle to late Miocene,
and therefore these physiographic features
formed afterward. Mordojovich (1981) previously considered this possibility and found it
problematic because no Neogene marine deposits occur in the highest summits (~1000 m) of
the Coastal Cordillera, which consist of metamorphic and plutonic basement rocks. However, flat erosional surfaces interpreted as relict
uplifted peneplains by Farías et al. (2008) occur
on the top of this mountain range as well as on
the summits of hill belts that transect the Longitudinal Valley. It is likely that these flat surfaces
are relatively young because the rainy climate
of this region is not conducive to their preservation. Supporting data for this interpretation
are the cosmogenic 10Be and 26Al age dates for
these low-relief elevated surfaces in the Coastal
Cordillera at 38°30′S, which suggest their
Pliocene to Holocene uplift (Rehak, 2008). In
addition, apatite fission-track dating of samples
collected in a vertical transect in the basement
of the Coastal Cordillera at 37°S–38°S suggest
an episode of accelerated denudation beginning
ca. 5 Ma (Glodny et al., 2008). It is therefore
conceivable that the Coastal Cordillera and the
transversal hill belts of the Longitudinal Valley
in this area were tectonically uplifted after deposition of Miocene marine strata, and those beds
pushed up to the highest elevations were obliterated by erosion.
CONCLUSIONS
Sedimentologic, ichnologic, and micropaleontologic studies of Neogene marine deposits in
the present Coastal Cordillera and Intermediate
Depression of south-central Chile reveal a major
event of subsidence that resulted in deep-marine
sedimentation (>2000 m) along the entire width
of the forearc during the middle to late Miocene. Subsidence of the margin during that period is attributed to subduction erosion, which
contradicts previous studies that interpreted the
nearly stable position of the magmatic arc since
the Jurassic as evidence that the current trench
line has remained stationary. The anomalously
thin crust that characterizes this area may have
facilitated forearc subsidence and allowed the
Miocene transgression to advance much farther
inland here than in other regions of Chile. It is
probable that the region of the present Coastal
Cordillera and Longitudinal Depression was entirely submerged during the Miocene and that
these physiographic features did not exist until
the Pliocene.
ACKNOWLEDGMENTS
This research was funded by Fondecyt Projects
3060051, 11080115, and 1110914 of Conicyt (Comisión Nacional de Investigación Científica y Tecnológica) and the IRD (Institut de Recherche pour le
Développement). Luis Buatois also received financial
support from Natural Sciences and Engineering Research Council (NSERC, Canada) Discovery grants
311726-05/08. We thank former IRD Director Gerard
Hérail, and Sergio Villagrán, driver of the institution
for their help. We are also grateful to Paul Duhart and
others at the regional office of Sernageomín (Servicio Nacional de Geología y Minería) for their logistics support and field assistance. ENAP (The Chilean
National Petroleum Company) kindly allowed us to
access their microfossil slides from wells drilled in the
study area. James Ackerson is acknowledged for allowing access to the Hacienda Parga.
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