Early Cenozoic uplift of the Puna Plateau, Central Andes, based on
stable isotope paleoaltimetry of hydrated volcanic glass
Robin R. Canavan1*, Barbara Carrapa2, Mark T. Clementz1, Jay Quade2, Peter G. DeCelles2, and Lindsay M.
Schoenbohm3
1
Department of Geology and Geophysics, University of Wyoming, Laramie, Wyoming 82010, USA
Department of Geosciences, University of Arizona, Tucson, Arizona 85721, USA
3
Department of Chemical and Physical Sciences, University of Toronto Mississauga, Mississauga, Ontario L5L 1C6, Canada
2
ABSTRACT
Uplift of the Central Andes is largely thought to have occurred during the past 10 m.y.
based on paleoaltimetry studies from the Altiplano of Bolivia. However, the spatio-temporal
uplift history may not be uniform across the Central Andes. We present new stable isotopic
results from the Salar de Antofalla, Salina del Fraile, and Arizaro Basin on the Puna Plateau
(24°–26°S) of northwestern Argentina. Samples of volcanic glass give δDpaleowater values and
modeled paleoelevations that indicate an elevated (~4 km) Puna Plateau since ca. 36 Ma with
limited (<1 km) changes in elevation since then and, unlike the Altiplano, no evidence of largemagnitude uplift during the Miocene.
INTRODUCTION
Today the Central Andes constitute the largest barrier to zonal moisture transport on Earth,
having a profound effect on South American
precipitation patterns (Strecker et al., 2007).
Their uplift would have had a dramatic impact
on regional climate and faunal and floral diversification (Ehlers and Poulsen, 2009; Hoorn et
al., 2010). Thus, an outstanding issue not only in
the field of Andean geodynamics but also in the
fields of climate sciences, biology, and ecology
is the timing of uplift of the Andes. From a geological perspective, understanding the timing of
uplift could help in evaluating the importance of
a variety of mountain-building processes, such as
arc magmatic activity (Kay et al., 1994), crustal
thickening in the retro-arc region (Kley and Monaldi, 1998; McQuarrie et al., 2005), and cyclical
lithospheric removal (DeCelles et al., 2009).
The timing and mechanism of uplift of the
Central Andes remains hotly debated. Stable
isotope and paleobotanical methods of paleoaltimetry used in the Altiplano (15°–20°S, Fig. 1)
suggest a 2.5–3 km increase in elevation between ca. 10.3 and 6.8 Ma as a result of wholesale removal of mantle lithosphere (GregoryWodzicki, 2000; Garzione et al., 2008). Isotopic
evidence from the Bolivian Eastern Cordillera
(Leier et al., 2013) finds more spatial and temporal (between 24 Ma and 15 Ma) variability
in the history of elevation change. Isotopic evidence of seasonal precipitation during the Miocene in the Subandean foreland is used as evidence of the deflection of the low-level jet due to
uplift of the Eastern Cordillera between 10 Ma
and 8 Ma (Mulch et al., 2010). In contrast, climate models (Ehlers and Poulsen, 2009) have
*Current address: Department of Geology and
Geophysics, Yale University, New Haven, Connecticut 06511, USA.
been used to argue for a more gradual uplift of
the Andes since 25 Ma. A major impediment to
understanding the paleoelevation history of the
Central Andes is the lack of data from the Puna
Plateau, which forms the southern half of the
Central Andes (24°–26°S, Fig. 1). We present a
stable isotope paleoaltimetry study of the Puna
Plateau, and demonstrate that elevations of
~4 km have persisted in this region since 36 Ma.
This observation has significant implications for
the causes of high elevation throughout the Central Andes.
GEOLOGICAL AND CLIMATIC
SETTING OF THE CENTRAL ANDES
The Central Andes include the Western Cordillera (modern arc), the Central Andean Plateau
(Altiplano and Puna Plateau), and the Eastern
Cordillera (Fig. 1). Moisture drawn from the
Amazon lowlands and Paraguayan Chaco Basin
produces intense rain-out (>2 m/yr) below 3 km
elevation along the eastern flank of the Andes,
and aridity westward into the Andean interior
(Strecker et al., 2007; Ehlers and Poulsen, 2009)
(Fig. 1). On average, the Puna Plateau is higher
(4400 m above sea level [masl]), drier, and more
rugged than the Altiplano (3800 masl) today
(Allmendinger et al., 1997; Strecker et al., 2007;
Garzione et al., 2008; Fig. 1). Despite these
modern differences, the two regions share similar deformation histories. Both the Altiplano
and Puna Plateau were occupied by a regional
foreland basin during the early Cenozoic, and
transformed into the modern Andean orogen
through crustal shortening and magmatic activity from the Paleogene onward (Sempere et al.,
1997; DeCelles and Horton, 2003; DeCelles et
al., 2011; Carrapa et al., 2011). Geophysical and
geochemical studies suggest that the mantle lithosphere beneath the Puna Plateau and Altiplano
is gravitationally unstable and subject to local
GEOLOGY, May 2014; v. 42; no. 5; p. 447–450; Data Repository item 2014157
|
doi:10.1130/G35239.1
|
or regional-scale foundering events (Kay et al.,
1994; Bianchi et al., 2013; Ducea et al., 2013).
Crustal thickness is highly variable (45–75 km)
under the Puna Plateau and is commonly thinner compared to that under the Altiplano (60–
80 km) (Yuan et al., 2002; Bianchi et al., 2013)
(Fig. 1). Thicker crust and greater shortening estimates for the Altiplano and Eastern Cordillera
of Bolivia (McQuarrie et al., 2005; Kley and
Monaldi, 1998) suggest that if shortening alone
were responsible for uplift, elevations should be
higher in the Altiplano. That this is not the case
suggests that other processes, such as underplating and lithospheric removal, may be controlling regional elevation differences between the
Puna Plateau and Altiplano.
STABLE ISOTOPE PALEOALTIMETRY
Stable isotope paleoaltimetry is based on the
observation that precipitation at higher elevations is progressively more depleted in both 18O
and deuterium (D) isotopes due to orographic
uplift and associated Rayleigh distillation (Rowley, 2007). To quantify paleoelevations on the
Puna Plateau, deuterium isotope ratios of hydrated volcanic glass (δDglass) from 33 volcanic
tuff and ignimbrite samples were determined
from the Salar de Antofalla (~3600 masl), Salina del Fraile (~3700 masl), and Arizaro Basin
(~3900 masl) (Fig. 1). The host deposits are
fluvial and eolian sandstones, and perennial and
ephemeral lacustrine shales with a total stratigraphic thickness of more than 1000 m in the
Salina del Fraile area and 3000 m in the Arizaro
Basin (Figs. DR1–DR8 in the GSA Data Repository1). New U-Pb zircon ages from 23 sampled
tuffs in the Salar de Antofalla and Salina del
Fraile areas range from 37.1 ± 8.4 Ma to 0.5
± 0.5 Ma (Tables DR1 and DR2 in the Data Repository). Previous U-Pb zircon ages from tuffs
in the Arizaro Basin range from 34.8 ± 2.3 Ma
to 0.4 ± 1.0 Ma (Boyd, 2010). Combined, these
data provide a stable isotope paleoaltimetry record from late Eocene (ca. 37 Ma) to modern for
the Puna Plateau.
1
GSA Data Repository item 2014157, detailed
methods and discussion supported by 14 figures and
2 data tables, is available online at www.geosociety
.org/pubs/ft2014.htm, or on request from editing@
geosociety.org or Documents Secretary, GSA, P.O.
Box 9140, Boulder, CO 80301, USA.
Published online XX Month 2014
GEOLOGY
©
2014 Geological
2014 | of
www.gsapubs.org
America. For permission to copy, contact Copyright Permissions, GSA, or editing@geosociety.org.
| May Society
447
Figure 1. A: Geologic
70° W Sub 65° W
map of Salar de Arizaro,
67° W
66° W
an
68° W
de
Arizaro Basin, Salar de
an
Antofalla, and Salina del
N
N
o
r
Fraile region of Puna
a
20° S
iz
Plateau,
northwestern
C
Ar
e
Argentina. Localities for
d
r
hydrated volcanic glass
la
Salta 25° S
samples are depicted.
Sa
For locations of mea25° S D
Puna
sured stratigraphic sections where more than
Sierra
one sample was colPamp s
eanas
lected, section number is
le
30° S
indicated. These include
ai
r
lF
sections in Salina del
de
Precipitation (mm/year)
Fraile: 2SF (n = 5), 6–7SF
a
500 km
<100
lin 2 SF
>2,000
(n = 4), 3–4, 8–9 SF (n = 4);
a
26° S
S
and in Salar de Antofalla:
4
6-7 SF
2LQ (n = 2) and 1SJ (n =
2 LQ 1 SJ
2
Altiplano Plateau
Arizaro Basin (n=10)
2). B: Shaded relief map
crust
with mean annual preSalina del Fraile (n=14)
3-4, 8-9 SF
-100
cipitation and major tecmantle lithosphere
Salar de Antofalla (n=9)
50
km
tonomorphic provinces,
-200
750
500
250
modified after Strecker
et al. (2007). S.B. Ranges
4
Undifferentiated
Cretaceous sedimentary rocks
east of the Puna Plateau
2
Active volcanic arc
Puna Plateau
Paleozoicmetasedimentary
igneous rocks
are the Santa Barbara
Paleozoic
Upper Cenozoic-Quaternary
crust
Ranges. Black box derocks
Paleo. metasedimentary rocks
salar deposits
mantle lithosphere
-100
notes area mapped in A.
Precambrian basement rocks
Cenozoic
sedimentary
rocks
500
250
White lines indicate locations of cross sections in
C and D. C,D: Cross sections showing topography and subsurface structural differences between Altiplano and Puna plateaus (GregoryWodzicki, 2000; Allmendinger et al., 1997; Kay et al., 1994; Yuan et al., 2002; Bianchi et al., 2013). Units are in kilometers; note change in
vertical scale at sea level.
A
B
rn Cord
Easte
o
lan
Sa
nges
lar
de
An
tof
all
a
Pre-Cordillera
S.B. R
a
ordillera
illera
Western
C
Ranges
ip
Alt
C
D
448
-130
Figure 2. U-Pb zircon ages
(Ma) and δDglass (‰, Vienna
standard mean ocean
water, VSMOW) values of
glass from sampled tuffs
and ignimbrites (y-axis is
reversed). Error bars for
age and δD are 2σ. Gray
panel shows modern
(younger than 0.55 Ma) average and range of δDglass
of samples from Dettinger
(2013). Asterisk denotes
samples thought to be
hydrated by evaporated
surface waters.
-120
δDglass (VSMOW, ‰)
δD of Volcanic Glass
Primary volcanic glass contains 0.1–0.3 wt%
magmatic water. Within 1000–10,000 yr after
eruption, it becomes hydrated by meteoric water until it reaches ~5 wt% water (Friedman et
al., 1993; Cassel et al., 2012). During hydration,
hydrolysis and proton (H+) exchange results
in a bonded silica “gel” layer which seals the
isotopic composition of initial hydration waters
(Cailleteau et al., 2008; Mulch et al., 2008; Cassel et al., 2012; Dettinger, 2013).
We found a large range in δDglass values
(Fig. 2; Table DR1), from −127.7‰ ± 3.3‰ to
−67.9‰ ± 2.4‰. This spread in δDglass values is
greater than that found today (−125‰ to −74‰)
based on δDglass values from modern (<0.55 Ma)
samples (n = 7) collected on the Puna Plateau
from 2973 masl to 4757 masl (Dettinger and
Quade, 2014). This greater range in δDglass
values likely reflects varying contributions of
different source water (evaporatively enriched
surface waters, snow melt, or high-elevation
waters depleted in deuterium). Results from two
samples (Eastash 1 and Eastash 2) have especially high δDglass values, ~6‰ higher than the
upper range of modern glass from the Puna Plateau (Fig. 2). These values are likely the result
of glass hydration by evaporatively enriched
surface waters, producing underestimated elevations. These scenarios are plausible as Eastash 1 is from the upper Miocene Sijes Formation, which contains numerous evaporites, and
-110
-100
-90
-80
-70
Arizaro Basin
Salina del Fraile
Salar de Antofalla
* *
< 0.55
-60 Ma
0
10
20
30
40
50
Age (Ma)
Eastash 2 is from a bedded tuff in Quaternary
alluvium. A similar scenario may account for
the higher-than-expected δDglass values for
one young Arizaro Basin sample (ARB09-9,
−80.4‰ ± 2.2‰, 0.4 ± 1.0 Ma), which was also
collected from Quaternary alluvium, although
its δDglass value is still within the range of other
modern glasses (Fig. 2).
Water content is highly variable in our glass
samples with most falling between ~4 and 12
wt% (Fig. DR9). Samples 3SF4, 6SF664, and
9SF3.8 have H2O >10 wt% and similar H concentration to a kaolinite standard (Figs. DR9–
DR12), suggesting contamination by adhering
clay particles, and therefore were not used to interpret paleoelevations. The δDglass values were
converted into δDpaleowater (δDpw) values using
the empirically derived fractionation factor (α)
of 1.0343 published by Friedman et al. (1993).
This fractionation factor has been applied successfully in previous paleoelevation studies using hydrated volcanic glass (e.g., Mulch et al.,
www.gsapubs.org
|
May 2014
|
GEOLOGY
Modeled Paleoelevations
Predictions of paleoelevation are made using
an atmospheric thermodynamic model based
on a change in δ18Oprecipitation between low-altitude and high-altitude locations (isotopic lapse
rate), written as Δ(δ18Op) (Rowley, 2007). This
model has been used to reconstruct paleoelevations for the Sierra Nevada (California), Colorado Plateau, Altiplano, and Himalaya (Cassel
et al., 2012; Mix et al., 2011; Rowley and Garzione, 2007). Due to the large error inherent in
the model because of its theoretical basis, paleoelevation estimates should be interpreted as
a potential range as opposed to absolute elevation. The strength of the Rowley (2007) model
is that it can incorporate different temperatures
(T) and relative humidities (RH), making it
useful for time periods for which an isotopic
lapse rate is unknown. However, it is important to choose appropriate values of RH and
T to reduce error in the paleoelevation results.
Because a major hydrologic shift occurs at the
Eocene-Oligocene boundary as Earth moved
from a greenhouse climate toward the cooler
and drier climate of today, we used the Rowley
(2007) model with Eocene parameters (RH =
80% ± 3%, T = 300 K) modeled from Huber
and Caballero (2003). To estimate paleoelevations for Oligocene and younger samples, we
use the modern model from Rowley (2007)
which utilizes all possible modern values of
starting temperature and relative humidity for
Δ(δ18Op) between 0‰ and –25‰.
In order to generate the model’s Δ(δ18Op)
term, we followed the method of Mix et al.
(2011), which involves converting δDpw values to δ18Opw values using the global meteoric
water line (GMWL) (Craig, 1961). In an effort
to minimize error associated with this conversion, we modified the method by using a local
meteoric water line (LMWL: δDmw = 8δ18Omw
+ 12.6) for northwestern Argentina from Dettinger (2013). We opted to use the LMWL
for the Oligocene to recent samples and the
GMWL for the Eocene samples, because the
“icehouse” conditions of the Oligocene and
younger time periods are closer to today’s conditions than the “greenhouse” conditions of the
Eocene. In addition, model results suggest that
the Eocene GMWL was very similar to that of
today (Speelman et al., 2010).
To account for changing climate throughout
the Cenozoic, we use epoch-specific low-elevation modeled or fossil δ18O and δD values to
GEOLOGY
|
May 2014
|
www.gsapubs.org
generate the Rowley (2007) model’s Δ(δ18Op)
term. These low-elevation values were chosen
based on their proximity to the assumed source
of moisture, which today is the subtropical Atlantic (Ehlers and Poulsen, 2009). We assume
the same source in the past because South
America has maintained a relatively constant
latitude throughout the Cenozoic (Sdrolias and
Müller, 2006). See the Data Repository for further discussion of MWLs and moisture source.
For the Eocene, we used a modeled δDpw value of −20‰ from Speelman et al. (2010) as our
low-altitude input, which converted to a δ18Opw
value of –3.8‰ using the GMWL. Rodent fossils from <1 km elevation in Brazil provided
a δ18Opw value of −3‰ (Bershaw et al., 2010)
for the Oligocene, and a modeled δ18Opw value
of −2‰ was used for the Miocene (Poulsen et
al., 2010). For both the Pliocene and Quaternary samples we used an average δ18Opw value
of −2.4‰ from three coastal Brazilian Global
Network of Isotopes in Precipitation (GNIP)
stations (Belém, Salvador, and Rio de Janeiro;
IAEA/WMO, 2006). Any error in model results
due to using this modern average for Pliocene
samples is within the limits of uncertainty of the
model itself.
Modeled mean elevation estimates indicate
a Puna Plateau height of between 3.4 ± 1.7 km
and 5.3 ± 2.8 km in the late Eocene. For the most
part, these upper Eocene samples come from the
Salina del Fraile and Salar de Antofalla, with
only one sample from the Arizaro Basin, which
gives a much lower average paleoelevation estimate; however, given the ±2.3 m.y. error for this
sample’s U-Pb zircon age, it may be more appropriate to group it with the Oligocene sample
subset. Paleoelevation estimates are between 3.2
± 1.4 km and 3.7 ± 1.5 km in the Oligocene, up
to 4.7 ± 1.8 km in the Miocene, and 3.9 ± 1.6 km
to 4.1 ± 1.6 km in the Pliocene and Quaternary
(very near modern elevations) (Fig. 3).
DISCUSSION AND CONCLUSIONS
Paleoelevation estimates from the Puna Plateau presented here indicate high elevation
(~4 km) since at least 36 Ma. High elevations
could be supported by a thickened crust beneath
the Puna Plateau associated with shortening and
eastward migration of the orogenic strain front
into the westernmost Eastern Cordillera during
the late Eocene (Carrapa and DeCelles, 2008)
and by magmatic underplating which is thought
to have increased crustal thickness below the
Puna Plateau from ~30–35 km during the Jurassic to ~45 km by the late Eocene (Haschke et
al., 2002). Speculations of small-scale (<1 km)
uplift events are tenuous because they are within
error of the model, but nonetheless, they may
be significant as they correlate with other geological evidence for uplift. For example, the
Miocene (ca. 25–18 Ma) event (<500 m; Fig. 3)
corresponds to deformation and exhumation in
the Eastern Cordillera (Coutand et al., 2006;
Carrapa et al., 2011), which is consistent with
significant crustal thickening and uplift in the
region at this time.
Unlike the Bolivian Altiplano and Eastern
Cordillera, there is no evidence of large-magnitude uplift of the Puna Plateau during the Miocene (Garzione et al., 2008; Leier et al., 2013),
and small-scale lithospheric removal under the
Puna Plateau (Ducea et al., 2013) can only account for limited uplift (<1 km). It is clear from
our results that there are important along-strike
variations in uplift between the Puna Plateau
and Altiplano. In turn, this implies a complex
lithospheric structure under the Central Andes
as monitored by geophysical data (Yuan et al.,
2002; Bianchi et al., 2013). High elevation since
36 Ma of part of the Central Andes should be
Figure 3. Estimated pa7000
leoelevations for Puna
Plateau based on Row6000
ley (2007) model. Error
Modern
bars are 2σ. Vertical gray
Puna
shaded region shows avPlateau
5000
erage and range of elevations for modern sample
locations. Black dashed
4000
line connecting samples
is three-point running average curve for paleoel3000
evation, excluding samples discussed in text
2000
thought to be characterized by evaporatively enBolivian
riched surface waters or
Altiplano
1000
contaminated by clays.
Dark gray band shows
Plio
Miocene
paleoelevation estimates
0
for Bolivian Eastern Cor0
10
20
dillera from Leier et al.
(2013). Light gray band
Age (Ma)
shows paleoelevation estimates for Bolivian Altiplano from Garzione et al. (2008).
Paleoelevation (m)
2008; Mix et al., 2011; Cassel et al., 2012).
Using this fractionation factor, our δDpw values
range from −97.8‰ ± 3.4‰ to −35.9‰ ± 2.5‰,
which is a greater range in δD values than what
is found today (−75.1‰ to −40.3‰) in sampled
meteoric waters (Dettinger, 2013). See the Data
Repository for further discussion of weight percent H2O and the fractionation factor.
?
Eastern
Cordillera
Oligocene Eocene
30
40
449
incorporated into future models of Andean tectonics, geodynamics, and paleoclimate.
ACKNOWLEDGMENTS
This research was funded by the National Science
Foundation (NSF) grant EAR-0911577, the Geological Society of America grant 9668-11, and the University of Wyoming. Partial support from NSF was provided through the University of Arizona’s Laserchron
Facility and by Exxon-Mobil through the University
of Arizona. We acknowledge the staff of the International Atomic Energy Agency and contributing institutions for collecting GNIP data. We thank G. Gehrels,
M. Pecha, and the staff of both the Laserchron Facility
and the University of Wyoming’s Stable Isotope Facility, including C. Cook and C. MacDonald. In addition,
we thank R. Alonso, F. Rivera, J. Boyd, S. Kim, W.
White, and B. Foreman who helped during field and
laboratory work.
REFERENCES CITED
Allmendinger, R.W., Jordan, T.E., Kay, S.M., and
Isacks, B.L., 1997, The evolution of the Altiplano-Puna Plateau of the Central Andes: Annual
Review of Earth and Planetary Sciences, v. 25,
p. 139–174, doi:10.1146/annurev.earth.25.1.139.
Bershaw, J., Garzione, C.N., Higgins, P., MacFadden,
B.J., Anaya, F., and Alvarenga, H., 2010, Spatialtemporal changes in Andean plateau climate and
elevation from stable isotopes of mammal teeth:
Earth and Planetary Science Letters, v. 289,
p. 530–538, doi:10.1016/j.epsl.2009.11.047.
Bianchi, M., Heit, B., Jakovlev, A., Yuan, X., Kay, S.M.,
Sandvol, E., Alonso, R.N., Coira, B., Brown, L.,
Kind, R., and Comte, D., 2013, Teleseismic tomography of the southern Puna plateau in Argentina and adjacent regions: Tectonophysics, v. 586,
p. 65–83, doi:10.1016/j.tecto.2012.11.016.
Boyd, J., 2010, Tectonic evolution of the Arizaro Basin
of the Puna Plateau, NW Argentina: Implications
for plateau-scale processes [M.S. thesis]: Laramie, Wyoming, University of Wyoming, 83 p.
Cailleteau, C., Angeli, F., Devreux, F., Gin, S., Jestin, J., Jollivet, P., and Spalla, O., 2008, Insight
into silicate-glass corrosion mechanisms: Nature Materials, v. 7, p. 978–983, doi:10.1038/
nmat2301.
Carrapa, B., and DeCelles, P.G., 2008, Eocene exhumation and basin development in the Puna
of northwestern Argentina: Tectonics, v. 27,
TC1015, doi:10.1029/2007TC002127.
Carrapa, B., Trimble, J.D., and Stockli, D.F., 2011, Patterns and timing of exhumation and deformation
in the Eastern Cordillera of NW Argentina revealed by (U-Th)/He thermochronology: Tectonics, v. 30, TC3003, doi:10.1029/2010TC002707.
Cassel, E.J., Graham, S.A., Chamberlain, C.P., and
Henry, C.P., 2012, Early Cenozoic topography, morphology, and tectonics of the northern
Sierra Nevada and western Basin and Range:
Geosphere, v. 8, p. 229–249, doi:10.1130
/GES00671.1.
Coutand, I., Carrapa, B., Deeken, A., Schmitt, A., Sobel, E.R., and Strecker, M.R., 2006, Propagation of orographic barriers along an active range
front: Insights from sandstone petrography and
detrital apatite fission-track thermochronology in the intermontane Angastaco Basin, NW
Argentina: Basin Research, v. 18, p. 1–26,
doi:10.1111/j.1365-2117.2006.00283.x.
Craig, H., 1961, Isotopic variations in meteoric waters: Science, v. 133, p. 1702–1703, doi:10.1126
/science.133.3465.1702.
DeCelles, P., and Horton, B.K., 2003, Early to middle Tertiary foreland basin development and
450
the history of Andean crustal shortening in Bolivia: Geological Society of America Bulletin,
v. 115, p. 58–77, doi:10.1130/0016-7606(2003)
115<0058:ETMTFB>2.0.CO;2.
DeCelles, P.G., Ducea, M.N., Kapp, P., and Zandt,
G., 2009, Cyclicity in Cordilleran orogenic
systems: Nature Geoscience, v. 2, p. 251–257,
doi:10.1038/ngeo469.
DeCelles, P.G., Carrapa, B., Horton, B.K., and Gehrels, G.E., 2011, Cenozoic foreland basin system in the central Andes of northwestern Argentina: Implications for Andean geodynamics
and modes of deformation: Tectonics, v. 30,
TC6013, doi:10.1029/2011TC002948.
Dettinger, M.P., 2013, Calibrating and testing the volcanic glass paleoaltimeter in South America [M.S.
thesis]: Tucson, Arizona, University of Arizona,
53 p.
Ducea, M.N., Seclaman, A.C., Murray, K.E., Jianu,
D., and Schoenbohm, L.M., 2013, Mantle-drip
magmatism beneath the Altiplano-Puna plateau, central Andes: Geology, v. 41, p. 915–918,
doi:10.1130/G34509.1.
Ehlers, T.A., and Poulsen, C.J., 2009, Influence of
Andean uplift on climate and paleoaltimetry estimates: Earth and Planetary Science Letters, v. 281,
p. 238–248, doi:10.1016/j.epsl.2009.02.026.
Friedman, I., Gleason, J., and Warden, A., 1993, Ancient climate from deuterium content of water
in volcanic glass, in Swart, P.K., et al., eds.,
Climate change in continental isotopic records:
American Geophysical Union Geophysical
Monograph 78, p. 309–319.
Garzione, C.N., Hoke, G.D., Libarkin, J.C., Withers, S., MacFadden, B., Eiler, J., Ghosh, P., and
Mulch, A., 2008, Rise of the Andes: Science,
v. 320, p. 1304–1307, doi:10.1126/science
.1148615.
Gregory-Wodzicki, K.M., 2000, Uplift history of the
Central Andes: A review: Geological Society
of America Bulletin, v. 112, p. 1091–1105, doi:
10.1130/0016-7606(2000)112<1091:UHOTCA
>2.0.CO;2.
Haschke, M., Siebel, W., Günther, A., and Scheuber,
E., 2002, Repeated crustal thickening and recycling during the Andean orogeny in north Chile
(21°−26°S): Journal of Geophysical Research,
v. 170, doi:10.1029/2001JB000328.
Hoorn, C., and 17 others, 2010, Amazonia through
time: Andean uplift, climate change, landscape
evolution, and biodiversity: Science, v. 330,
p. 927–931, doi:10.1126/science.1194585.
Huber, M., and Caballero, R., 2003, Eocene El Niño:
Evidence for robust tropical dynamics in the
“hothouse”: Science, v. 299, p. 877–881, doi:
10.1126/science.1078766.
IAEA/WMO (International Atomic Energy Agency/
World Meteorological Organization), 2006,
Global Network of Isotopes in Precipitation:
The GNIP Database: http://www.iaea.org/water
(April 2013).
Kay, S.M., Coira, B., and Viramonte, J., 1994, Young
mafic back arc volcanic rocks as indicators
of continental lithospheric delamination beneath the Argentine Puna plateau, central Andes: Journal of Geophysical Research, v. 99,
p. 24,323–24,339, doi:10.1029/94JB00896.
Kley, J., and Monaldi, C.R., 1998, Tectonic shortening and crustal thickness in the Central Andes:
How good is the correlation?: Geology, v. 26,
p. 723–726, doi:10.1130/0091-7613(1998)026
<0723:TSACTI>2.3.CO;2.
Leier, A., McQuarrie, N., Garzione, C., and Eiler,
J., 2013, Stable isotope evidence for multiple
pulses of rapid surface uplift in the Central An-
des, Bolivia: Earth and Planetary Science Letters, v. 371–372, p. 49–58, doi:10.1016/j.epsl
.2013.04.025.
McQuarrie, N., Horton, B.K., Zandt, G., Beck, S.,
and DeCelles, P.G., 2005, Lithospheric evolution of the Andean fold–thrust belt, Bolivia,
and the origin of the central Andean plateau:
Tectonophysics, v. 399, p. 15–37, doi:10.1016/j
.tecto.2004.12.013.
Mix, H.T., Mulch, A., Kent-Corson, M.L., and Chamberlain, C.P., 2011, Cenozoic migration of topography in the North American Cordillera: Geology, v. 39, p. 87–90, doi:10.1130/G31450.1.
Mulch, A., Sarna-Wojcicki, A.M., Perkins, M.E.,
and Chamberlain, C.P., 2008, A Miocene to
Pleistocene climate and elevation record of the
Sierra Nevada (California): Proceedings of the
National Academy of Sciences of the United
States of America, v. 105, p. 6819–6824, doi:
10.1073/pnas.0708811105.
Mulch, A., Uba, C.E., Strecker, M.R., Schoenberg,
R., and Chamberlain, C.P., 2010, Late Miocene
climate variability and surface elevation in the
central Andes: Earth and Planetary Science
Letters, v. 290, p. 173–182, doi:10.1016/j.epsl
.2009.12.019.
Poulsen, C.J., Ehlers, T.A., and Insel, N., 2010, Onset
of convective rainfall during gradual late Miocene rise of the Central Andes: Science, v. 328,
p. 490–493, doi:10.1126/science.1185078.
Rowley, D.B., 2007, Stable isotope-based paleoaltimetery: Theory and validation, in Kohn, M.J.,
ed., Paleoaltimetry: Geochemical and thermodynamic approaches: Mineralogical Society
of America Reviews in Mineralogy and Geochemistry 66, p. 23–52.
Rowley, D.B., and Garzione, C.N., 2007, Stable isotope-based paleoaltimetry: Annual Review of
Earth and Planetary Sciences, v. 35, p. 463–508,
doi:10.1146/annurev.earth.35.031306.140155.
Sdrolias, M., and Müller, R.D., 2006, Controls on
back-arc basin formation: Geochemistry Geophysics Geosystems, v. 7, Q04016, doi:10.1029
/2005GC001090.
Sempere, T., Butler, R.F., Richards, D.R., Marshall,
L.G., Sharp, W., and Swisher, C.C., 1997, Stratigraphy and chronology of late Cretaceous–
early Paleogene strata in Bolivia and northern
Argentina: Geological Society of America Bulletin, v. 109, p. 709–727, doi:10.1130/0016-7606
(1997)109<0709:SACOUC>2.3.CO;2.
Speelman, E.N., Sewall, J.O., Noone, D., Huber,
M., von der Heydt, A., Sinninghe Damsté, J.,
and Reichart, G., 2010, Modeling the influence
of a reduced equator-to-pole sea surface temperature gradient on the distribution of water
isotopes in the Early/Middle Eocene: Earth and
Planetary Science Letters, v. 298, p. 57–65,
doi:10.1016/j.epsl.2010.07.026.
Strecker, M.R., Alonso, R.N., Bookhagen, B., Carrapa,
B., Hilley, G.E., Sobel, E.R., and Trauth, M.H.,
2007, Tectonics and climate of the southern central Andes: Annual Review of Earth and Planetary Sciences, v. 35, p. 747–787, doi:10.1146
/annurev.earth.35.031306.140158.
Yuan, X., Sobolev, S.V., and Kind, R., 2002, Moho
topography in the central Andes and its geodynamic implication: Earth and Planetary Science Letters, v. 199, p. 389–402, doi:10.1016
/S0012-821X(02)00589-7.
Manuscript received 28 October 2013
Revised manuscript received 10 February 2014
Manuscript accepted 28 February 2014
Printed in USA
www.gsapubs.org
|
May 2014
|
GEOLOGY