Late Quaternary stratigraphy, sedimentology and plateau (northwest Argentina)

Basin Research (2013) 25, 638–658, doi: 10.1111/bre.12025
Late Quaternary stratigraphy, sedimentology and
geochemistry of an underfilled lake basin in the Puna
plateau (northwest Argentina)
Michael M. McGlue,* Andrew S. Cohen,† Geoffrey S. Ellis* and Andrew L. Kowler†
*Central Energy Resources Science Center, U. S. Geological Survey, Denver, CO, USA
†Department of Geosciences, The University of Arizona, Tucson, AZ, USA
Depositional models of ancient lakes in thin-skinned retroarc foreland basins rarely benefit from
appropriate Quaternary analogues. To address this, we present new stratigraphic, sedimentological
and geochemical analyses of four radiocarbon-dated sediment cores from the Pozuelos Basin (PB;
northwest Argentina) that capture the evolution of this low-accommodation Puna basin over the past
ca. 43 cal kyr. Strata from the PB are interpreted as accumulations of a highly variable, underfilled
lake system represented by lake-plain/littoral, profundal, palustrine, saline lake and playa facies
associations. The vertical stacking of facies is asymmetric, with transgressive and thin organic-rich
highstand deposits underlying thicker, organic-poor regressive deposits. The major controls on
depositional architecture and basin palaeogeography are tectonics and climate. Accommodation
space was derived from piggyback basin-forming flexural subsidence and Miocene-Quaternary normal faulting associated with incorporation of the basin into the Andean hinterland. Sediment and
water supply was modulated by variability in the South American summer monsoon, and perennial
lake deposits correlate in time with several well-known late Pleistocene wet periods on the Altiplano/Puna plateau. Our results shed new light on lake expansion–contraction dynamics in the PB
in particular and provide a deeper understanding of Puna basin lakes in general.
Modern lakes occur in a wide variety of tectonic settings,
and sediments recovered from such basins prove valuable
in geological and palaeoenvironmental research. Unlike
lakes formed by glacial or fluvial processes, tectonic lakes
typically persist on the landscape for 104 years, often
producing thick depositional sequences that can archive
climatic, biological, and surficial processes with high resolution (Olsen, 1990; Colman et al., 1995; GierlowskiKordesch & Park, 2004; McGlue et al., 2008). Despite
many decades of study, major gaps exist in our understanding of several types of modern tectonic lakes, particularly those associated with retroarc foreland basin
systems (DeCelles & Guiles, 1996). Lake formation is relatively well understood in thick-skinned forelands, and
data concerning ancient lakes exist in great abundance for
these basins (e.g. Eocene Green River Formation of western North America; Eugster & Hardie, 1975; Smith
et al., 2003). In this setting, lakes may form as erosionresistant basement blocks rising along steep reverse faults
Correspondence: Michael M. McGlue, Central Energy
Resources Science Center, U. S. Geological Survey, P.O. Box
25046, M.S. 977, Denver, CO 80225, USA. E-mail: [email protected]
that cause loading, flexure, and sediment-starved depressions to develop (e.g. Carroll et al., 2006).
The formation of lakes in thin-skinned forelands, however, is more complicated. In these orogens, topographic
closure in the proximal foredeep is hindered by erosion of
the thrust belt, as high rates of sediment accumulation
( 10 1 mm year 1; Sinha & Friend, 1994) balance or
overwhelm available accommodation space. Accordingly,
lakes are scarce in these settings and usually exist only
when the watershed geology is carbonate-rich, thereby
favouring rivers with low ratios of bedload to dissolved
load (e.g. Drummond et al., 1996; Zaleha, 2006). In contrast, lake formation is more likely in the hinterland
regions of a thin-skinned foreland system, as these higher
and drier environments may lack the ability to transport
significant sediment loads. Climate is critical to topographic closure and lake type in these intracontinental settings, due to its effect on sediment and water supply,
which helps to govern interactions between lake level and
the basin sill (Carroll & Bohacs, 1999). This is especially
true along mountain fronts such as the Andes, where rising air masses lose much moisture at low elevations.
Indeed, lakes and wetlands are conspicuous components of high-altitude basins in the thin-skinned central
Andes, providing vital habitat for a wide range of endemic
species and a key resource base for local human populations (Caziani et al., 2001; N
~ez et al., 2002). Yet, from
© 2013 The Authors
Basin Research © 2013 John Wiley & Sons Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists
Puna basin lacustrine deposystems
the perspective of basin analysis, many of these modern
deposystems are understudied. This knowledge deficit
limits the full use of lacustrine deposits in explorations of
ancient tectonic and climatic change in retroarc foreland
To address this gap, we studied four sediment cores
from Laguna de los Pozuelos (LP) (Gangui, 1998). This
playa-lake occupies the centre of the Pozuelos Basin (PB),
a piggyback basin in the Puna plateau of northwest Argentina (Fig. 1). Radiocarbon-dated cores from LP provide
an excellent opportunity to characterize the stratigraphy,
sedimentology, and geochemistry of an underexplored
class of thin-skinned retroarc foreland basin lakes. Furthermore, sediments from LP provide the chance to assess
climate change in the Puna and its northern equivalent,
the Altiplano. Notably, palaeoclimate proxy records are
spatially complex and sometimes conflicting from the high
Andean plateau (Betancourt et al., 2000; Bobst et al.,
2001; Latorre et al., 2002; Godfrey et al., 2003; Fritz
et al., 2004; Chepstow-Lusty et al., 2005; Maldonado
et al., 2005; Placzek et al., 2006; Nester et al., 2007;
Quade et al., 2008; Blard et al., 2011). Prior actualistic
analyses of LP sediments (McGlue et al., 2012a) were
used to guide interpretations of these cores, and we present herein new insights into the facies architecture
and palaeogeography of the PB from ca. 43 cal ka BP to
The PB is an NNE-oriented, elongate (ca. 2750 km2) piggyback basin at ca. 22°S, 66°W. West-verging thrust
sheets carrying siliciclastic and volcanic Ordovician rocks
bound the flat-floored basin (Fig. 1a). Relief between the
basin floor (ca. 3663 m a.s.l.) and flanking ranges exceeds
450 m, but the basin spill point is <40 m above the modern playa-lake (Fig. 1b). Seismic stratigraphic analysis
and regional correlations suggest that PB formation and
synorogenic sedimentation began in the Oligocene, with
maximum subsidence occurring near the eastern-margin
thrust faults (Gangui, 1998). The basin is tectonically
complex and the most recent deformation is associated
with normal faulting and volcanism (Cladouhos et al.,
1994). Neogene ignimbrites are widespread along the
basin’s eastern flank, whereas small exposures of Cretaceous nonmarine sediments crop out near the southern
end of the basin. Miocene nonmarine carbonates, evaporites, and tuff (i.e. Cara Cara Formation; Cladouhos et al.,
1994) cropout along the eastern basin margin (Fig. 1a).
Total precipitation in the PB is ca. 320 mm year 1 and
monthly mean air temperatures range between 3 and
13°C (Legates & Willmott, 1990a, b). Rainfall, derived
almost entirely from eastern sources, is strongly seasonal,
with about 70% of the yearly total occurring during the
austral summer. Climate in the region is governed by the
Fig. 1. (a) Simplified geological map and cross-section of the Pozuelos Basin (PB). Cross-section location is marked by a dashed line.
Extensional lineaments (dotted lines) are from Caffe et al. (2002). LP, Laguna de los Pozuelos. SR, Sierra de Rinconada. SC, Sierra de
Cochinoca. SQ, Sierra de Quichagua. T, Tertiary. M, Miocene. O, Ordovician. K, Cretaceous. Q, Quaternary. GP, Group. (b) Shuttle
Radar Topography Mission digital elevation model of the PB illustrating the elongate basin shape and its spill-point, located ca. 35 m
above the basin floor. (c) Approximate location of the sediment cores discussed in the text, referenced to a recent shoreline of LP. Inset
map shows the position of the basin in South America.
© 2013 The Authors
Basin Research © 2013 John Wiley & Sons Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists
M.M. McGlue et al.
South American summer monsoon (SASM; Zhou & Lau,
1998). As with most of the Puna plateau, the El Ni~
noSouthern Oscillation (ENSO) phenomenon and North
Atlantic sea surface temperatures modulate patterns of
precipitation over the PB (Garreaud et al., 2009). Rainfall, topography, and soil moisture gradients control vegetation in the basin, which is a mixture of C3 and C4
grasses, shrubs, succulents, and macrophytes (Bonaventura et al., 1995; McGlue et al., 2012a).
Laguna de los Pozuelos is hydrologically closed, making it sensitive to changes in effective precipitation (P-E;
Teller & Last, 1990). The surface area of LP fluctuates
annually, and several lines of evidence indicate that it can
exceed ca. 135 km2 during years with above average precipitation (e.g. Mirande & Tracanna, 2009). Conversely,
intervals of prolonged drought commonly lead to its desiccation. The playa-lake floor is flat when LP is filled and
the maximum water depth is ca. 1.5 m. Weak southeasterly summer winds and stronger westerly winter winds
prevent LP from developing persistent stratification. LP
is fed by both groundwater and a small surface water
drainage network (Igarzabal, 1978). The Rıos Cincel and
Chico are more permanent and form small axial deltas at
the southern end of LP, whereas the Rıo Santa Catalina is
ephemeral and forms a seasonally subaerial, terminal splay
complex north of LP (Fig. 1c; McGlue et al., 2012a).
Numerous ephemeral streams form alluvial fans along
LP’s lateral margins.
Sediment cores were collected from the PB using a modified split-spoon sampler attached to a gasoline-powered
hammering device (Fig. 1c and Table S1). PVC liners
allowed incremental core sections to be collected by repeat
drives into the open borehole. Recovery per drive varied
(30–100%), most likely in response to vertical changes in
water content and sediment density. Individual core sections were sealed in the field, then shipped to LacCore
(University of Minnesota) and subsequently analysed for
physical properties using a GEOTEK multi-sensor scanner. Magnetic susceptibility, gamma-ray attenuation,
lithostratigraphic markers, and radiocarbon data were
used to: (1) differentiate between intact stratigraphy and
sediments collapsed from the sides of the borehole; (2)
vertically correlate the intact stratigraphy; and (3) create
composite stratigraphic logs for each borehole.
Facies analysis was conducted on freshly split core surfaces. Particle sizes were estimated using a grain size card,
and sedimentary components were assessed using a combination of smear slides, ca. 125-lm sieved residues, and
powder X-ray diffraction. Shortly after core splitting, discrete sediment samples (2–3 cm3) were collected every ca.
12–15 cm and freeze–dried prior to further analysis. Elemental and stable isotopic analyses of sedimentary organic
matter (OM) were conducted at the University of Arizona
(UA) to provide insights into biomass production, preser-
vation, and provenance. Total organic carbon (TOC), total
nitrogen and d13COM were measured on a Costechâ
(Costech Analytical Technologies Inc., Valencia, CA,
USA)elemental analyser coupled to a continuous-flow gasratio mass spectrometer (Finnigan Delta PlusXLâ;
Thermo Fisher Scientific Inc., Waltham, MA, USA). To
remove carbonate minerals that could influence isotope
values, samples were pretreated at room temperature for
several hours using 1M HCl, then washed four times in de
ionized water and air dried. Samples were combusted in
the elemental analyser. Standardization was based on acetanilide for elemental concentration, NBS-22 and USGS24 for d13COM. Precision was better than 0.09 for
d13COM based on repeated internal standards. Atomic C/
N ratios were corrected for contributions of inorganic
nitrogen following the procedure outlined by Talbot
(2001). Total inorganic carbon (TIC) and biogenic silica
(BiSi) analyses were conducted at LacCore. Weight percent of TIC was determined using a UIC Inc.â (Joliet, IL,
USA) total carbon coulometer and provided quantitative
constraints on carbonate content. Analytical precision
associated with this analysis was ca. 0.20%. BiSi analyses
utilized multiple extractions of hot alkaline digestions following a modified protocol designed by Demaster (1979).
Reported values had an analytical precision of 1.0%.
Rock-Eval pyrolysis was conducted on select de-calcified
samples (n = 20) from deep intervals in the cores at the University of Houston, to help discriminate the source of OM
and infer environmental conditions during deposition (Espitalie et al., 1977; Katz, 1983).
Due to the importance of age control for interpreting
stratigraphy and palaeoclimate, the geochronology of
Quaternary lake sediments from the Puna and Altiplano
has been the subject of much research (Geyh et al.,
1999; Sylvestre et al., 1999). Potential pitfalls associated
with the radiocarbon dating of different organic materials in Andean lakes were summarized in detail by Geyh
et al. (1999) and Placzek et al. (2006). Briefly, samples
may be affected by the introduction of 14C-depleted
water into lake surface waters from various carbon reservoirs, producing old apparent ages. Alternatively,
sources of carbon in equilibrium with the atmosphere at
any given time can contaminate older sediments, resulting in spuriously young apparent ages.
As with many lakes in high-altitude arid catchments,
terrestrial organic materials are exceedingly rare in LP.
Thus, the initial approach for establishing the chronology
at LP focused on dating <63-lm sediment OM, primarily
using a low-temperature combustion technique (McGeehin et al., 2001). This procedure minimizes the potential
for older, clay-bound carbon to influence the age determination. Moreover, samples were exposed to an acid-baseacid pretreatment to remove any carbonate or humic acid
that could complicate age determinations. In all, 23 sediment samples were dated using this method (Table S2). In
addition, we dated seeds of the macrophyte Ruppia
(n = 4), which grow in LP today and are also preserved in
sediment cores. Carbon extraction and 14C analysis were
© 2013 The Authors
Basin Research © 2013 John Wiley & Sons Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists
Puna basin lacustrine deposystems
Fig. 2. Age models with error envelopes (light gray) for cores 2A, 3A, 4A and 6A. Organic matter dates used in the interpolation
appear as black diamonds, whereas excluded dates appear as white diamonds. Ruppia dates appear as gray diamonds. Inset maps show
the location of cores with respect to a recent shoreline of Laguna de los Pozuelos. Note slower sedimentation rates between ca. 19 and
3 cal ka BP, interpreted as evidence of maximum lowstand conditions.
carried out at the UA – Accelerator Mass Spectrometer
Facility. Conversion of 14C ages to median calendar ages
utilized the INTCAL09 calibration curve and the program
CALIB 6.0 (; Reimer et al.,
2009). Sediment samples that returned ‘post-bomb’ 14C
ages were calibrated using the Southern Hemisphere routine (Hua & Barbetti, 2004) in the program CALIBomb
( An age-depth
model was developed for each of the sediment cores using
a simple linear interpolation (Fig. 2).
Radiocarbon geochronology
A seed collected from a living Ruppia plant in 2006
returned a post-bomb age, which calibrated to 1959–1961
or 1983–1986 CE (Table S2). These dates constrain reservoir effects within modern LP to <50 years, because
aquatic floras assimilate only dissolved inorganic carbon
(DIC) and not carbon from the atmospheric reservoir.
The water and sediment chemistry of LP suggests that
discharge from deep aquifers makes negligible contributions to the playa-lake’s hydrologic balance (McGlue
et al., 2012a). Furthermore, minimal exchange of groundwater DIC with 14C-dead inorganic carbon from Miocene
carbonates on the eastern PB margin is likely in the modern system (Fig. 1a). Therefore, we conclude that at present, DIC is most strongly influenced by dilute runoff and
shallow groundwater sources.
If deep palaeolakes occupied the PB and experienced
prolonged intervals of water column stratification, then
the DIC in these palaeoenvironments could have been out
of equilibrium with atmospheric CO2. A similar situation
would occur if greater than modern contributions from
C-depleted groundwater flowed into the basin. To
assess the extent of these potential 14C-reservoir effects,
we employed paired dating on palaeo-Ruppia and coevally
© 2013 The Authors
Basin Research © 2013 John Wiley & Sons Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists
M.M. McGlue et al.
deposited sedimentary OM. A single Ruppia seed from
core 2A (depth 434 cm) returned a 14C age younger than
that of coeval OM, suggesting either a reworked origin of
the OM or a reservoir effect influencing the apparent age
of the seed. In contrast, OM from depth 526.5 cm yielded
a date several thousand years younger than Ruppia seeds
from depths 540 and 434 cm (Table S2). We interpret the
OM age at depth 526.5 cm as spurious, most likely due to
laboratory contamination; it was excluded from the age
model. As reversals do not occur in the other cores
(Fig. 2), we favour an age model for core 2A that excludes
the two anomalies, employing the Ruppia date at 434 cm
in place of its paired OM counterpart. Additional support
for this interpretation is provided by the concordance of
core 2A’s resultant chronostratigraphy with those derived
for the other cores. Because our data do not permit quantification of possible reservoir effects within individual
stratigraphic units, we do not apply a reservoir correction
and report calibrated ages with the caveat that these values may represent maximum ages.
Our oldest record (ca. 43 cal ka BP) comes from core
3A, whereas the lowermost dated samples from the other
cores are slightly younger (ca. 37–38 cal ka BP; Fig. 2).
The ages of core tops across the PB are highly variable,
ranging from 1966 CE to ca. 1 cal ka BP. We interpret
nonmodern ages to indicate eolian deflation and bioturbation of the basin floor, which may have been subaerially
exposed at different times in the late Holocene. Supporting text on radiocarbon-derived sedimentation rates is
located in the online archive (Fig. S1).
Facies analysis
Thirteen facies types were recognized in the analysis
(Fig. 3 and Table S3). These facies were laterally continuous and could be identified in all of the cores across the
nearly 5-km span of the study site, except where noted.
Facies A – massive clay
This facies consisted of ungraded, massive, red-brown
clay with up to ca. 20% dispersed quartz and mica silt
(Fig. 3a). Mudcracks were present at the tops of beds.
Minor abundances of ostracodes, diatoms, and inorganic
carbonates (low-Mg calcite and aragonite) were also present. The average TOC, TIC, and BiSi concentrations
were 0.8, 0.5, and 2.3 wt.% respectively. Beds were ca.
10–35 cm thick and they exhibited nonerosional, indistinct basal contacts (transition over 1 cm).
Facies A is attributed to suspension sedimentation in the
extant playa-lake, LP (Igarzabal, 1978; McGlue et al.,
2012a). Playa waters were well-oxygenated due to wind
mixing, which accounts for oxidized sediment colours
and poor OM preservation, as reflected in low TOC content (Cohen, 2003). Siliciclastic grains were carried to the
playa-lake by sheet floods and wind (Hardie et al., 1978),
but laminations have been disrupted by bioturbation
(McGlue et al., 2012a). Polygonal cracks are common on
the PB floor today, and observed mudcracks are attributed
to desiccation. Although the sediments of many playalakes consist of evaporites, mixed siliciclastic-carbonate
playas similar to LP may develop where runoff and standing water permeate into deeper aquifers (e.g. Turnbridge,
1984; Chivas et al., 1986). Alternatively, flooding may
dissolve soluble evaporites on a seasonal basis (Smoot &
Lowenstein, 1991). Diatoms, particularly Cocconeis placentula, Nitzschia hungarica and Navicula sp., represented
the dominant source of BiSi within this facies (Maidana
et al., 1998).
Facies B – weakly stratified clayey sand
This facies consisted of ungraded, crudely bedded (cmscale) to massive, red-brown, very fine- to mediumgrained clayey sands with rare vertical cracks (Fig. 3a).
Diatoms were present in minor abundances. The average
TOC, TIC and BiSi concentrations were 1.4, 0.2 and
2.2 wt.% respectively. Framework grains were moderate
to well-sorted and consisted of quartz, muscovite, and
plagioclase feldspar. Beds were ca. 15–20 cm thick and
exhibited sharp (change over <1 mm), weakly erosive
basal contacts.
Facies B is attributed to subaerial sheetfloods on the
extant Rıo Santa Catalina terminal splay complex
(McGlue et al., 2012a). Stratification was produced by
upper flow regime, plane bed conditions during sand
deposition (Fedo & Cooper, 1990; Horton & Schmitt,
1996). Waning (decelerating) flows of flood events probably provided the source of clays (Fisher et al., 2008). Terminal splay sedimentation occurs where unconfined
floods enter a closed basin, but do not significantly raise
playa-lake levels (e.g. Lang et al., 2004; Fisher et al.,
2008). Vertical cracks are attributed to subaerial exposure
and desiccation of the terminal splay complex. The headwaters of the Rıo Santa Catalina reside in the Sierra de
Rinconada (western PB margin; Fig. 1a), which means
that detrital grains were recycled from Ordovician marine
sandstones and shale.
Facies C – massive oxide-rich silty clay with
silt laminae and irregular silt pods
This facies consisted of ungraded, massive, tan silty clay
with Fe-oxide mottles (Fig. 3b). Tilted or flat laminations
(0.5–3 mm thick) or thin (1–3 mm) jagged–edged pods of
silt were scattered throughout (Fig. 3b). Mudcracks were
rare and filled with fine sand or mud. Whole and fragmented ostracodes were common, and XRD scans
detected traces of halite. The average TOC, TIC and BiSi
concentrations were 0.4, 0.4 and 2.8 wt.% respectively.
© 2013 The Authors
Basin Research © 2013 John Wiley & Sons Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists
Puna basin lacustrine deposystems
Fig. 3. Sediment core photographs from the Pozuelos Basin, illustrating the major facies encountered in the study area. Cores are ca.
3.5 cm wide and tops are to the right. (a) Massive red clay (Facies A) and weakly stratified clayey sand (Facies B) of the playa facies
association (Unit I). Facies B bedding, most likely produced by sheetfloods, has been disrupted by the coring process. (b) Massive
oxide-rich tan clay (Facies C) of the saline lake facies association (top Unit II), with arrows marking the location of tilted and flat laminae. This facies is interpreted as the maximum lowstand in the stratigraphic framework. (c) Massive black pyrite-rich mud (Facies D)
of the saline lake facies association (basal Unit II), with arrow marking a burrow-like oxidation feature. (d) Massive green silty clay
(Facies E) of the palustrine facies association (Unit III), with arrow marking the location of dispersed pebbles. (e) Laminated diatom
ooze (Facies F1), showing an example of the sharp contact with underlying transgressive deposits. (F) Laminated diatom ooze (Facies
F2). Arrow marks thick laminations comprised of calcified Chara debris and ostracodes. (g) Laminated to thinly bedded OM-rich silty
clay (Facies G), a sublittoral deposit from core 2A. These sediments are coeval to Facies F, suggesting shoaling in the direction of the
core site. (h) Mottled clays (Facies H), with crudely laminated and disrupted beds of macrophyte debris marked by an arrow.
(i) Thinly interbedded green sand and clay (Facies I2; marked by arrow), most likely produced by sheetfloods. (j) Normally graded and
massive sands from core 3A. (k) Normally graded, matrix-supported gravel and coarse sand (Facies M) overlying Facies L from core
Beds were ca. 45–90 cm thick and exhibited diffuse to
indistinct (transition over 1 mm to 3–4 cm) contacts that
lacked erosion.
Facies C is attributed to deposition and reworking in a
dry mudflat. Colour, Fe-oxides, mudcracks and low TOC
all point towards very shallow water, oxidizing conditions,
mostly likely with prolonged intervals of subaerial expo-
sure and desiccation of the basin floor (Plummer & Gostin, 1981; Smoot, 1983; Demicco & GierlowskiKordesch, 1986). Silt laminae were comprised of rounded
quartz and highly refractory micas that suggest eolian
processes helped to shape this environment (e.g. Keen &
Shane, 1990; Rosen, 1994). The paucity of fine sedimentary structures, massive bedding, and jagged-edged silt
pods are interpreted as evidence that bioturbation has
altered primary depositional textures. This type of
reworking, primarily accomplished by waterbirds, is
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M.M. McGlue et al.
common on LP’s fringing mudflats today (McGlue et al.,
2012a). Alternatively, these pods may have accumulated
in the cavities of thin efflorescent salt crusts, which is consistent with the presence of halite.
Facies D – massive pyrite-rich mud
This facies consisted of ungraded, massive, black pyrite-rich mud with minor diatoms (Fig. 3c). Rare calcite laminae (1–2 mm thick) were also encountered.
Pyrite occurred as fine framboidal aggregates, in some
cases filling the interior cavities of pennate diatoms.
The average TOC, TIC and BiSi concentrations were
1.0, 0.9 and 2.4 wt.% respectively. Dolomite and
halite were detected on XRD scans in this facies. Beds
were ca. 40–75 cm thick and basal contacts were
Facies D is attributed to deposition by suspension fallout in a perennial saline lake. Authigenic calcite laminae suggest a lake whose bottom periodically escaped
reworking by waves and bioturbation. Facies D shares
a number of similarities with the Pleistocene saline
lake facies of the Badwater Basin (Death Valley, California, United States), including (1) rapid (hours to
days) oxidation of black muds to a gray-green colour;
(2) the presence of burrow-like oxidation features
(Fig. 3c); (3) carbonate laminae; and (4) low average
TOC (Roberts et al., 1994). Low TOC may be the
result of bacterial oxidation of OM in the presence of
sulphate, which also helps to explain the abundance of
pyrite (Potter et al., 2005). Evaporative loss of lake
water is indicated by the presence of dolomite and
halite, but the dominance of detrital sediment suggests
that this lake was sustained by surface inflows (e.g.
Smoot & Lowenstein, 1991; Rosen, 1994).
Facies E – massive silty clay
This facies consisted of ungraded, massive, green silty
clay. Ostracodes, diatoms, calcified Chara stems, and disseminated macrophyte debris were minor to abundant in
Facies E. Rare outsized (2–3 cm) sedimentary rock clasts
were encountered in this facies at core site 2A (Fig. 3d).
The average TOC, TIC and BiSi concentrations were
0.8, 0.8 and 3.4 wt.% respectively. Beds were ca.
65–250 cm thick and exhibited diffuse to indistinct (transition between ca. 1 and 40 mm) basal contacts that lacked
Facies E is attributed to deposition by suspension fallout
in the littoral zone of a perennial lake. This interpretation
is supported by the presence of massive bedding and calcareous microbenthos, which are common features in the
strata of shallow, well-oxygenated lakes marked by low
gradient floors (Galloway & Hobday, 1996; Blair &
McPherson, 2008). Formation of carbonate stem casts follows CO2 uptake by Chara during photosynthesis, which
requires shallow water in a lake’s photic zone (Anadon
et al., 2002). Dispersed pebbles near the base of Facies E
are interpreted to reflect rare deposition by gravity flows
during storms, as alternative mechanisms like ice or biological rafting (Bennett et al., 1996) were much less likely
in this environment.
Facies F – laminated diatom ooze
This facies consisted of thinly or thickly laminated
(up to 5–6 mm), dark green-brown diatom ooze
(Fig. 3e, f). Two variants of this facies were observed.
Facies F1 consisted of thinly laminated ooze that was
characterized by high diatom diversity (Fig. 3e). Facies
F2 consisted of ooze with intermittent thick laminations of calcified charophyte debris, ostracodes, and
pyrite-encrusted macrophyte fragments (Fig. 3f). The
average TOC, TIC and BiSi concentrations were 2.3,
0.7 and 5.3 wt.% respectively. Traces of halite were
also detected on XRD scans. Beds were typically
20–45 cm thick and were encountered in the axis of
the basin (cores 4A, 3A and 6A). Basal contacts were
planar and sharp.
The variants of Facies F are attributed to deposition by
suspension fallout in the profundal zone of a perennial
lake. Broadly similar diatom-rich laminites have been
identified as open lacustrine (below the photic zone)
deposits in a number of other intermontane basins (Fritz
et al., 2004; Rigsby et al., 2005). Facies F1 was defined by
fine planar laminations and high TOC, which suggest a
low-energy profundal zone that escaped wave reworking.
Highly reduced sediment colours point to oxygen-deficient waters that may have been influenced by seasonal
water-column stratification (e.g. Katz, 1995). We interpret deltaic processes to have influenced the development
of Facies F2, with wave action assisting in the concentration of biogenic debris.
Facies G – laminated to thinly bedded OMrich silty clay
This facies consisted of laminated to thinly bedded (up to
5 cm), brown and green silty clay (Fig. 3g). Fragments of
aquatic macrophytes (Ruppia sp. and Najas sp.) and Chara were the principal constituents of brown laminae,
whereas green laminae and thin beds consisted of silty
clay with diatoms. The average TOC, TIC, and BiSi concentrations were 2.1, 1.2 and 1.5 wt.% respectively. Beds
ranged from ca. 24 to 70 cm thick and were exclusively
found in core 2A, on the western basin margin. Basal contacts were planar and sharp.
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Puna basin lacustrine deposystems
Facies G is attributed to deposition by suspension fallout in the sub littoral zone of a perennial lake. Preservation of laminations suggests a low-energy lake floor,
possibly influenced by wave action but deep enough to
escape flamingo bioturbation (upper limit of ca.
50 cm; Mascitti & Casta~
nera, 2006). Abundant silt and
macrophyte fragments, as well as the relatively high
TIC content, imply a contrast in water depth between
Facies G and the laterally equivalent Facies F, which
is interpreted as a shoaling of the lake floor towards
the location of core 2A (Fig. 1c).
Facies H – mottled clays
This facies consisted of massive, mottled dark greenolive-brown clays with variably abundant macrophyte,
diatom, and calcified Chara debris (Fig. 3h). In some
cases, biogenic and inorganic components were crudely
interbedded, but disrupted fabrics were most common.
The average concentrations of TOC, TIC and BiSi were
1.7, 0.8, and 4.6 wt.% respectively. Beds in cores 2A and
6A were ca. 44–70 cm thick and basal contacts lacked erosion and were diffuse to indistinct (transition over 1 mm
to 3–4 cm).
Facies H is attributed to deposition by suspension fallout
at the margin of a perennial lake. Mottling suggests that
the environment was marked by variable water saturation
and interaction of iron with oxygenated or reduced pore
waters (Freytet & Verrecchia, 2002; Lindbo et al., 2010).
Although clear evidence of pedogenesis is absent, mottles
may reflect root traces and the macrophyte remains could
signify the presence of a supra-littoral wetland (e.g. Liutkus & Ashley, 2003). Elevated concentrations of BiSi
and TOC suggest relatively high productivity for this
environment. Lake-margin wetlands identified in the rock
record often preserve diatom-rich sediments, due to the
prevalence of organic acids and low pH levels during
deposition (Deocampo & Ashley, 1999).
Facies I – thinly interbedded sand and silty
Two variants of Facies I were encountered in this study.
Facies I1 consisted of thinly interbedded packages of dark
brown sand (0.5–2 cm) and green silty clay with diatoms.
Facies I1 exhibited very fine- to medium-grained sands
and wavy bedding, and was present at the base of Core
6A. Facies I2, found at the base of Core 4A, exhibited dark
green, very fine- to coarse-grained sands interbedded with
silty clay. Sedimentary structures were absent from Facies
I2. XRD scans determined that the framework grains were
comprised of quartz, volcanic rock fragments (dominantly
plagioclase feldspar) and muscovite. The average concen-
trations of TOC, TIC and BiSi were 0.2, 0.5 and
1.9 wt.% respectively. Beds were <10 cm thick and the
nature of the contacts with underlying units is unknown.
Facies I1 is attributed to deposition by sheetfloods associated with the transition to a subaqueous delta, due to the
interfingering of sand with diatom-rich silty clays. In this
case, sedimentation was likely rapid as distal sheetfloods
became inundated by rapidly rising lake levels at a low
point in the basin (Smoot, 1985). Facies I2 is also attributed to deposition by sheetfloods in a more proximal subaerial distributary environment (Hampton & Horton,
2007). Interbedded sand and muds are common deposits
that form as unconfined floods spread out over alluvial
plains (Gierlowski-Kordesch & Rust, 1994). The abundance of volcanic rock fragments in Facies I2 accounts for
the sub-equal percentages of quartz and andesine detected
by XRD. The most likely source of these sediments was
erosion of the eastern PB margin, where Ordovician and
Miocene volcanic arc lithologies are exposed in the Sierra
de Cochinoca (Fig. 1a).
Facies J – massive sands
This facies consisted of massive, ungraded to poorly
inversely graded, green, moderate to well sorted, fine- to
coarse-grained sands with a few granule clasts encountered near the top of the bed (Fig. 3j). The basal contact
was planar and weakly erosional. The bed of Facies J was
ca. 42 cm thick and present in core 3A.
Facies J is attributed to deposition by hyperconcentrated
flows or sheetfloods associated with a subaerial deltaic
environment. Rapid deposition of sand from turbulent
suspension precluded the development of bedforms
(Smith, 1986; Horton & Schmitt, 1996). Facies J overlies
interbedded matrix-and clast-supported gravels, which
may reflect deposition from the dilute, waning stage of
floods (Pierson & Costa, 1987). Smoot (1983) documented
similar, massive-bedded sheetflood sandstones associated
with a closed-basin lake in the western USA. The coarsefraction mineralogy of Facies J is similar to Facies I2,
suggesting a common provenance within the Sierra de
Facies K – normally graded sands
This facies consisted of normally graded, unstratified,
green, poor to well-sorted fine to very coarse sands with
dispersed gravel typical near the base (Fig. 3j). Sedimentary structures were absent. Beds were ca. 15–25 cm thick
and present in cores 3A and 4A. Basal contacts were irregular and erosional.
© 2013 The Authors
Basin Research © 2013 John Wiley & Sons Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists
M.M. McGlue et al.
Facies K is attributed to deposition by hyperconcentrated
flows or sheetfloods associated with a subaerial deltaic
environment. In the case of channelized hyperconcentrated stream flows, sedimentation from turbulent suspension
produced the normal grading, but dispersive pressure
and buoyancy generated by the high clast concentration
prevented bedform development (Smith, 1986). Deposits
similar to Facies K are also described for sheetfloods,
where grading and poor sorting reflect decelerating, highconcentration unconfined flows (Smoot, 1983; Smoot &
Lowenstein, 1991).The abundance of volcanic rock fragments accounts for the green colour and elevated plagioclase (andesine) content of these sands.
Facies L – normally graded clast-supported
This facies consisted of normally graded, green, poor to
moderately sorted, clast-supported very fine to fine gravel
(Fig. 3k). Outsized coarse pebble clasts (>2 cm) were
found near the bases of these deposits. A muddy sand
matrix accounted for 10–35% of these deposits and distribution normal grading (both clasts and matrix fine
upward) was common. XRD scans indicated that quartz,
plagioclase (andesine in cores 3A and 4A; albite in core
2A), and muscovite were the dominant framework grains
in Facies L. Beds were ca. 5–60 cm thick and basal contacts were planar and nonerosive.
Facies L is attributed to deposition by hyperconcentrated
flows or sheetfloods associated with a subaerial deltaic
environment. Sediment was transported in hyperconcentrated flows by a combination of turbulence, grain dispersive pressure and buoyancy (Smith, 1986). Such flows
may have originated from torrential floods that incorporated water and sediment along their paths (Sohn et al.,
1999). Although traction-produced sedimentary structures were not apparent, the normal grading and matrix
abundance may have been produced by clast interactions
in a turbulent cohesionless flow (Horton & Schmitt,
1996). Facies L in cores 3A and 4A have similar framework mineralogy to Facies I, J and K; in these cores, this
facies is likely to be made up of sediments derived from
Sierra de Cochinoca (Fig. 1a). In contrast, Facies L in
core 2A appears to be derived from the Sierra de Rinconada, based on the presence of Na-feldspars.
Facies M – normally graded matrixsupported gravels
This facies consisted of normally graded, green, poor to
moderately sorted, matrix-supported coarse sand to fine
gravel that lacked internal stratification (Fig. 3k). Coarsetail normal grading was typical in Facies M. A muddy
matrix accounted for >60% of the deposit. Beds of in core
3A were ca. 3–15 cm thick and basal contacts were nonerosive and planar to sub-planar (tilted).
Facies M is attributed to deposition by subaerial pseudoplastic debris flows associated with a subaerial deltaic or
alluvial fan environment. The cohesion of the finegrained matrix was the primary grain-support mechanism, but buoyancy and elevated pore pressure played a
role in supporting larger clasts (e.g. Lowe, 1982). Very
coarse-grained, thickly bedded, matrix-supported conglomerates often represent onshore debris flow lobes associated with lacustrine fan deltas (McPherson et al., 1987;
Blair & McPherson, 2008). However, several authors have
documented fine-grained, thinly bedded matrix-supported conglomerates in fan deltas, believed to be derived
from dilute debris flows with low matrix strength and dispersive pressure (Schultz, 1984; Horton & Schmitt,
Lithostratigraphy and facies associations
Five lithostratigraphic units (V–I, old to young) are present in the PB cores (Figs. 4 and 5). The vertical succession of facies associations is: lake-plain/littoral (Unit V);
profundal (Unit IV); palustrine (Unit III); saline lake
(Unit II); and playa (Unit I). Units V through II are interpreted as a lake expansion–contraction cycle, similar to
those that have been described for underfilled lake basins
(e.g. Pietras & Carroll, 2006). This cycle occurred from
ca. 43 to 3 cal ka BP and was deposited under variable P-E
conditions. Unit I represents the inception of a new cycle
that began after ca. 3 cal ka BP and was deposited under
negative P-E conditions, similar to the modern climate.
Lake-plain/littoral association (Unit V)
This association marks an interval of basin flooding and
transgression that began prior to ca. 43 cal ka BP. Its
thickness is ca. 5 to 135 cm and includes Facies I, J, K, L
and M. These facies are interpreted to represent a delta
that entered the PB along its eastern margin. Beds thin
and sediments become finer grained towards the core 6A
site, near the centre of the PB (Fig. 1c). Most of the sedimentological characteristics of Unit V indicate deposition
in the onshore portion of the delta by sheetfloods, hyperconcentrated stream flows and debris flows. The general
absence of interbedded fine-grained lake sediments supports this interpretation (Fig. 5; Horton & Schmitt,
1996). By contrast, interbedded sand and diatom-rich
silty clay (Facies I1) reflects distal sheetflood deposition
interacting with rising lake waters (Fig. 4b). The interpretation of delta type is constrained by the spatial distribution and number (n = 4) of cores in the dataset, which
does not permit a thorough evaluation of delta architecture (i.e. sheets versus channels). Additionally, the narrow
© 2013 The Authors
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Puna basin lacustrine deposystems
Fig. 4. (a) Stratigraphy of core 2A. (b) Stratigraphy of core 6A. 14C, radiocarbon age, presented as median calendar year before present. C, clay. Si, silt. Sa, sand. G, gravel. TIC, total inorganic carbon. BiSi, biogenic silica. TOC, total organic carbon.
width of the cores could preclude the full expression of
structures like climbing ripples, which clearly indicate
decelerating flows and are typical in sheetflooding environments.
Sheet, ‘Gilbert-type’, birdfoot-type or fan deltas are
commonly associated with perennial lakes in closed
basins, and in some cases, deltas with combined characteristics have been documented (Smoot & Lowenstein,
1991). Most available data support either a sheet or fan
delta interpretation for Unit V. McPherson et al. (1987)
explained that coarse-grained fan deltas usually exhibit
deposits from both gravity flows and sheetfloods, which is
consistent with Facies M, I2, J and K. However, Smoot &
Lowenstein (1991) noted that identifying fan deltas in
closed basins is problematic because of the potential for
rapid lake level fluctuations and wave reworking of subaerial alluvial fan sediments. Evidence of shoreline deposits are absent in Unit V, which suggest that either the
onshore–offshore transition was not captured in the cores,
or that beach deposits were poorly developed. Poorly
developed shorelines are characteristic of basins with
gently sloping, uniform floors where sheet deltas are
© 2013 The Authors
Basin Research © 2013 John Wiley & Sons Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists
M.M. McGlue et al.
Fig. 5. (a) Stratigraphy of core 3A. (b) Stratigraphy of core 4A. 14C, radiocarbon age, presented as median calendar year before present. C, clay. Si, silt. Sa, sand. G, gravel. TIC, total inorganic carbon. BiSi, biogenic silica. TOC, total organic carbon.
common (Gierlowski-Kordesch & Rust, 1994). Thinly
bedded, graded sands and gravels (Facies J and L respectively) are common where unconfined sheetfloods enter
closed basins (Smoot, 1983). Regardless of delta type,
Unit V deposits clearly indicate significant surface water
inflows to the PB, which suggest more positive P-E than
Sediment mineralogy suggests that a lateral river(s)
draining the Ca-plagioclase-rich, andesitic and dacitic
lavas of the Sierra de Cochinoca fed this ancient delta
(Fig. 1a; Caffe et al., 2002). Facies L encountered at the
base of core 2A is also interpreted as a delta, which
entered the PB on its western margin, with sediment supply from the Sierra de Rinconada (Fig. 1a). Here, the
presence of well-preserved Ruppia seeds and pyrite suggests a shallow subaqueous depositional environment.
Profundal association (Unit IV)
This association marks a short-lived episode of lake-level
highstand from ca. 43–36 cal ka BP. Its thickness is ca. 25
to 120 cm and includes Facies E, F1/F2, and G. Profundal association deposits are separated from underlying
Unit V beds by sharp contacts, suggesting rapid lateral
lake expansion. An offshore, relatively deep lacustrine
environment was interpreted based on the presence of
fine-grained laminated ooze. The southerly deepening
bathymetric trend implied by the lake-plain/littoral facies
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Puna basin lacustrine deposystems
association is borne out by the presence of profundal
Facies F2 at the base of core 6A. This facies suggests that
in addition to suspension settling, wave activity influenced sediment accumulation in this palaeolake. Timeequivalent, interbedded Facies G and Facies E at the core
2A site are interpreted as a sublittoral environment of
deposition, situated several km west of the deepest zone
of the palaeolake (Fig. 4a). The presence of thick laminations and thin beds in Facies G provides additional support for shoaling towards the western PB margin, as
variations in laminae texture are known to accompany
changes in water depth in other lake systems (Smoot,
Expansion of a relatively large palaeolake in the PB
may have promoted water column stratification and
reducing conditions on the lake floor, which is consistent
with the preservation of fine laminations in Facies F1 and
elevated TOC in most of the Unit IV facies types (Katz,
1995; Cohen, 2003). Widespread bottom-water anoxia
appears unlikely for this palaeolake, however, due to the
abundance of in situ benthic invertebrate fossils in Facies
E, F2 and G. In concert with carbonates, traces of halite
identified in profundal oozes suggest that the lake’s
hydrology was closed during the deposition of Unit IV.
The presence of a large palaeolake suggests more positive
P-E than modern during the deposition of Unit IV.
Palustrine association (Unit III)
This association is interpreted to reflect shoreline retreat
from ca. 37–23 cal ka BP. Its thickness ranges from ca. 70
to 260 cm and includes Facies E and H. Palustrine association deposits are separated from the underlying Unit IV
beds by gradational contacts, suggesting gradual contraction of the lake under climate conditions marked by
declining P-E. The palustrine environment encompasses
both permanently and seasonally inundated areas (Freytet
& Verrecchia, 2002; Pietras & Carroll, 2006). In low
accommodation, low gradient deposystems like the PB,
the palustrine environment is highly sensitive to fluctuations in P-E, as minor changes in inflows can expose broad
areas of the basin floor.
A palustrine origin for Facies E is implied by abundant
ostracodes and macrophyte OM, which indicate the presence of a shallow, well-oxygenated, photic zone that was
conducive to benthic organisms and plant growth. Massive bedding features in Facies E were likely produced by
bioturbation and wave mixing. Adjacent lake plain environments likely supplied the ‘floating’ pebbles (Fig. 3d)
to Facies E during intense storms that produced gravity
flows that entered the lake. A shoreline wetland is indicated by the mottled clays (Facies H) and massive green
silty clays (Facies E) in cores 6A and 2A. Facies H, which
is characterized by mottling, thick laminations, and crudely bedded macrophyte debris, is similar to sediments
that have been linked to fluctuating, shallow lake environments in other Andean basins (Rigsby et al., 2005).
Saline lake association (Unit II)
This association marks a continued regressive phase from
ca. 23 to 3 cal ka BP. Its thickness ranges from ca. 95 to
180 cm and includes Facies C and D. Saline lake association deposits are separated from the underlying Unit III
beds by indistinct contacts. The depositional environments represented in this association include a perennial
saline lake and dry mudflats, suggesting that evaporation
(declining and negative P-E) played an increasingly
important role in Unit II time.
Basal sediments in Unit II relate to the presence of a
perennial saline lake and mark the final appearance of a
palaeolake in the PB (ca. 26–19 cal ka BP). We interpret
that this lake was smaller and held more saline waters than
the palaeolake represented by the underlying profundal
and palustrine facies association (Units IV and III). Nevertheless, the massive pyrite-rich silty clays of Facies D
are characteristic of a saline lake where relatively fresh
surface waters were important to basin hydrology and
sedimentation (Smoot & Lowenstein, 1991). Authigenic
calcite precipitation was common in this saline lake, as
revealed by discrete laminae and peaks in TIC (Figs. 4
and 5). By contrast, most of the carbonate of the profundal and palustrine facies associations is derived from
ostracodes and Chara. Halite and dolomite were detected
in XRD scans, but not on smear slides, which suggest to
us that palaeolake waters were relatively dilute. This condition is not unusual, however, as perennial saline lakes
are known to exist for thousands of years without precipitating evaporites (Smoot & Lowenstein, 1991; Roberts
et al., 1994)
The transition from Facies D to Facies C marks a
major palaeoenvironmental change in Unit II time. Maximum lowstand conditions in the PB are inferred from
Facies C, which exhibits the slowest sedimentation rates
in the record (Fig. 2). The most likely depositional environment for Facies C is a dry mudflat. This mudflat
developed due to the desiccation of the perennial saline
lake (Facies D) as climate became more arid and surface
and groundwater inflows to the basin were diminished.
Facies C exhibits desiccation cracks, tilted silt laminae,
and very low TOC, which are features that are consistent
with subaerial exposure and deflation of lake beds (Smoot
& Lowenstein, 1991). Traces of halite are likewise present
in Facies C, and may have formed by evaporative pumping of groundwater along the mudflat margins.
Playa association (Unit I)
This association marks the inception of a transgression
and new lake cycle in the PB that began after ca. 3 cal ka
BP. Its thickness ranges from ca. 16 to 54 cm and includes
Facies A and B. Playa association deposits are separated
from the underlying saline lake facies association by sharp
to indistinct contacts, suggesting an abrupt expansion of
the lake surface. Depositional environments represented
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M.M. McGlue et al.
in Unit I include the extant playa-lake and terminal splay
complex of the Rıo Santa Catalina. Effective precipitation
is interpreted to be similar to modern for Unit I, which
strongly contrasts with the climate that accompanied the
transgression of Unit V.
Today, the Rıo Santa Catalina maintains an axial flow
path in the PB (Fig. 1c), and the fine-grained deposits of
its terminal splay (interbedded Facies A and B) reflect the
geology of the western basin margin. This depositional
configuration is much different from the coarse-grained
delta of Unit V, whose mineralogy and colour indicate
palaeoflow from the eastern basin margin. Today, summer rainfall produces sheetfloods on the Rıo Santa Catalina (Facies B) that rarely interact with the extant playalake, whereas flooding and P-E is interpreted to have been
much higher during Unit V time. The playa-lake was
described in detail elsewhere (McGlue et al., 2012a), but
the massive clays of Facies A share a number of characteristics with disrupted playa mudstones observed in other
closed basins (Turnbridge, 1984; Gore, 1989).
Organic matter geochemistry
Bohacs et al. (2000) explained that basin type strongly
controls organic enrichment in lacustrine rocks, which is
a function of OM production, destruction, and dilution.
For the late Quaternary sediments of the PB, we examined bulk OM geochemistry (Fig. 6a, b) and Rock Eval
datasets (Fig. 6c), to assess the importance of each of
these major controls.
Production in lake basins refers to the primary productivity by plants, which can grow within the lake itself
(autochthonous) or in the lake’s watershed (allochthonous), requiring transport by wind or water prior to sedimentation. Commonly cited controls on productivity in
lake basins include solar energy, water chemistry and
nutrient availability (Passey et al., 2010). The production
of OM is relatively high in Unit IV, where profundal
facies average 2.3 wt.% TOC (Fig. 6a and Table S3).
Hydrogen index (HI) values for Facies F and G (profundal and adjacent sub littoral deposits) range from 92 to
287 mg HC g 1 TOC (Fig. 6c). In concert with microscopic observations, these Rock Eval data suggest autochthonous production from a mixture of algae and
macrophytes (i.e. Type II kerogens) in a late Pleistocene
lake that was likely larger than modern LP (Fig. 6c). The
enriched d13COM values for Unit IV (Fig. 6b) are attributed to the DIC-pool of the lake water, which was probably influenced by the kinetic effects of evaporation and
algal utilization of HCO3 as a carbon source for photosynthesis (Meyers & Teranes, 2001). The absence of
higher plant remains on smear slides suggests that transported terrestrial vegetation contributed very little to the
sedimentary OM of Unit IV, which is consistent with
observations from other underfilled lakes (Bohacs et al.,
2000). Mottled clays (Facies H) of the palustrine association also produce TOC peaks (Fig. 4) and mean HI values
of 240 mg HC g 1 TOC (Fig. 6c). Facies H creates a dis-
Fig. 6. Bulk organic matter (OM) geochemistry from the Pozuelos Basin. (a) Total organic carbon (TOC) – total organic nitrogen (TON) crossplot. Note elevated TOC concentrations of
profundal Facies F. (b) Carbon to nitrogen ratio (C/N) – carbon
isotope (d13COM) crossplot. Note broad C/N and enriched
d13COM values, interpreted as reflecting destructive processes
due to shallow bathymetry in all palaeolake environments. Note
that symbols are the same for panel A. (c) Rock Eval modified
van Krevelen diagram. HI, hydrogen index. OI, oxygen index.
Poor preservation of OM supports the interpretation of basin
underfilling in the late Quaternary.
tinct cloud on the left side of the C/N – d13COM crossplot, mostly likely signifying high diatom productivity in
the wetland environment (Fig. 6b). In the saline lake and
playa facies associations, accumulations of OM are generally low (Fig. 6a) and production appears to be dominated
by macrophytes, which are known to be isotopically
enriched in the PB (Fig. 6b; McGlue et al., 2012a).
Destruction refers to removal of OM prior to or during
sedimentation, typically by inorganic oxidation,
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Puna basin lacustrine deposystems
photo-oxidation, microbial respiration or ingestion by
metazoans (Bohacs et al., 2000). Destructive processes
are interpreted to be the most significant influence on
organic sedimentation in the low accommodation PB.
This is especially true late in Unit II, as very low sedimentation rates and frequent subaerial exposure limited
the preservation of OM in the dry mudflat environment
(average TOC = 0.4 wt.%). During Unit I time, the
playa-lake’s shallow bathymetry and polymixis conspired
to limit the preservation of mixed algal/macrophyte OM
(McGlue et al., 2012a). The C/N – d13COM crossplot
and relatively high oxygen index values indicate that oxidation may have affected the larger palaeolakes represented by Unit IV, III, and II strata (Fig. 6b, c). For
example, the diagenetic loss of labile nitrogen could help
produce the broad range of C/N values exhibited by these
facies associations, as terrestrial vegetation makes very
limited contributions to the sedimentary OM (Meyers &
Teranes, 2001). The heavy impact of oxidation is attributed to the relatively shallow bathymetry and limited
accommodation space inferred for all of the PleistoceneHolocene palaeolake environments in the PB.
Dilution refers to the reduction in the concentration of
OM in sediments due to deposition of siliciclastic or biogenic (e.g. diatom silica or shell carbonate) materials. Siliciclastic dilution is not a common process in the
depocentres of underfilled lake basins, as lake expansion
during highstands constrains transport and deposition of
coarse siliciclastic detritus to the basin margins (Smoot,
1983; Pietras & Carroll, 2006). Relatively high TOC values in laminated diatom oozes (Facies F) argues against
significant biogenic dilution during highstands. Dilution
is most apparent in our cores at the muddy tops of Unit V
deltaic deposits. Organic carbon concentrations for these
sediments (Facies I, J and L) never exceed 0.3 wt%
(Figs 4 and 5).
Depositional history and palaeogeography
Several lines of evidence suggest that during the past ca.
43 cal kyr, water levels in the PB remained below the
spill-point elevation, resulting in basin underfilling
(Fig. 7). Underfilled lake basins form where rates of
accommodation continually exceed the rate of sediment
plus water fill (Carroll & Bohacs, 1999). The cyclic strata
of the Wilkins Peak Formation (Green River Basin; Pietras & Carroll, 2006), which represent lake expansion–contraction dynamics, are perhaps the best studied
underfilled lake deposits. Quaternary PB strata share a
number of similarities with these Wilkins Peak cycles,
including (1) asymmetric facies stacking patterns, with
thin transgressive facies overlain by thicker regressive
intervals (Pietras et al., 2003); (2) scale, with cycles up to
ca. 5–6 m thick (Pietras & Carroll, 2006); (3) abundant
sedimentary evidence of evaporation and desiccation
(Carroll & Bohacs, 1999); (4) low overall TOC, with peak
organic enrichment just above the shoreline/transgressive
package (Bohacs et al., 2000); and (5) strongly coupled
accumulation rates and lake levels, such that highstand
deposits display the highest sedimentation rates (Bohacs
et al., 2000). Pietras & Carroll (2006) pointed out that the
vertical stacking of disparate facies in Wilkins Peak cycles
do not readily conform to Walther’s law of facies succession, making a strict sequence stratigraphic interpretation
a challenge. The low sedimentation rates, silt laminae,
low TOC concentrations, and diagenetically altered OM
exhibited by the dry mudflat deposits of Unit II (Facies
C) provide clear evidence of subaerial exposure and erosion; this is the best candidate for a sequence boundary in
the PB record.
Notably, Type II OM is present in the PB, which contrasts with the Type I OM that more commonly accumulates in underfilled basins (Fig. 6c). This is believed to be
linked to the dearth of accommodation space in the basin,
which led all palaeolakes to remain relatively shallow.
Whereas many underfilled basins are hypersaline, the saline-alkaline waters of palaeolakes in the PB allowed both
macrophytes and diatoms to flourish. This interpretation
is also supported by sediment mineralogy, as alkaline
earth carbonates are present throughout Units IV-I,
whereas halite is only present in traces. These data suggest that groundwater discharge from saline aquifers did
not control hydrology in the PB during the late Quaternary. Rather, fluctuating P-E and surface water inflows
appear to have been most important to palaeolake hydrochemistry.
Today, the landscape of the Puna plateau is marked by
internally drained basins, dry valleys, ephemeral playalakes, and salt flats, which reflect the arid climate. Precipitation at the PB falls in the austral summer, and due to its
latitude, Amazonian sources should play a key role in the
moisture balance (Garreaud et al., 2009). Our palaeogeographical reconstruction posits that palaeolakes larger
than modern LP occupied the basin during the late Pleistocene (Fig. 7). To sustain large lakes, most researchers
have suggested that large swings in P-E are necessary
(Hastenrath & Kutzbach, 1985). Others favour variability
in air temperature and glacial meltwater controls on lake
levels, which modelling has shown may be valid in certain
locales on the southern Altiplano (e.g. Blodgett et al.,
Several decades of research suggest that during the late
Pleistocene, a series of palaeolakes existed on the Altiplano (<250 km northwest of the PB; Fig. S2). The existence of palaeolakes Tauca (ca. 18–14 cal ka BP) and
Coipasa (ca. 13–11 cal ka BP) is almost universally agreed
upon, as data for these lakes exist in both drill core stratigraphy (Baker et al., 2001a, b) and shoreline records (Placzek et al., 2006). The presence, extent and timing of other
palaeolakes are the topic of ongoing debate. Placzek et al.
(2006) argued for a deep paleolake, Ouki, from ca.
© 2013 The Authors
Basin Research © 2013 John Wiley & Sons Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists
M.M. McGlue et al.
Fig. 7. Palaeogeographical sketch maps of the Pozuelos Basin (PB) from ca. 43 ka – present illustrating pervasive basin underfilling.
Lakes appear blue (darker hues signify greater relative water depths), and the basin landscape appears green (positive P-E) or brown
(negative P-E). (a) A shallow playa-lake with intermittent axial streams and fringing mudflat environments has existed since ca.
3 cal ka BP. Note that the Rıo Santa Catalina forms a fine-grained terminal splay near the northern end of the playa-lake. (b) A dry
mudflat, formed by subaerial exposure and eolian deflation of lake beds, likely during the early-middle Holocene. If Tauca and Coipasa
aged palaeolakes existed in the PB, evidence of them is missing. (c) A saline lake occupied the PB from ca. 26 to 19 cal ka BP, which
broadly correlates with the Sajsi lake cycle (Placzek et al., 2006; Blard et al., 2011). (d) From ca. 37 to 23 cal ka BP, a shallow paleolake
occupied the basin and produced thick regressive deposits in our cores. (e) A deeper paleolake existed in the PB from ca. 43 to
37 cal ka BP. Based on the present age model, maximum lake expansion correlates in time with the Minchin highstand and Heinrich
event 4, identified in other records from the Altiplano (Baker et al., 2001b; Fritz et al., 2004; Kanner et al., 2012). (f) Stacked coarse
siliciclastic units indicate the presence of a delta in the PB prior to ca. 43 cal ka BP. The mineralogy of deltaic deposits suggests palaeoflow from the east.
120–98 cal ka BP and several shallow palaeolakes, Salinas
(ca. 95–80 cal ka BP), Inca Huasi (ca. 46 cal ka BP), and
Sajsi (ca. 24–21 cal ka BP), on the basis of shoreline stratigraphy, radiocarbon and U/Th data. More recently,
Blard et al. (2011) reported that the Sajsi palaeolake
was moderately deep and refined its chronology to ca.
25–19 cal ka BP, coincident with the global Last Glacial
Maximum. By contrast, interpretations of drill core stratigraphy have focused on the deep Tauca and an older
lake known as Minchin, which is believed to have occupied the Uyuni Basin from ca. 46 to 36 cal ka BP (Baker
et al., 2001b; Fritz et al., 2004; Chepstow-Lusty et al.,
2005). The Sajsi palaeolake is not typically differentiated
in drill core strata and the age assigned to the Tauca palaeolake spans ca. 26–15 cal ka BP, whereas evidence of the
large Ouki palaeolake and indeed, large lakes in general
are absent prior to ca. 50 cal ka BP (Fritz et al., 2004,
© 2013 The Authors
Basin Research © 2013 John Wiley & Sons Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists
Puna basin lacustrine deposystems
A number of climatic hypotheses have been advanced
to explain rainfall variability on the Altiplano/Puna plateau. Some researchers favour variability in Southern
Hemisphere insolation and its effect on the SASM as the
dominant control on lake hydrology (Baker et al., 2001a,
b; Fritz et al., 2004; Hanselman et al., 2011). This group
focuses on the correlation among maximum summertime
insolation, increased interhemispheric meridional SST
gradients, and evidence of lake expansion on the Altiplano. North Atlantic SST gradients in particular are
considered critical to moisture in the Amazon, the source
region for precipitation in the central Andes (Baker et al.,
2001b). For example, cold SST anomalies and Heinrich
events (massive iceberg discharges prompting weak thermohaline circulation; Bond et al., 1992) in the Atlantic
are correlated with evidence of wet conditions at Lake
Titicaca and the Salar de Uyuni (Baker et al., 2001a; Fritz
et al., 2004; Blard et al., 2011). Additionally, a well-dated
speleothem record from Peru provides compelling evidence for linkages between Heinrich events and the intensity of the SASM over the past ca. 50 cal kyr (Kanner
et al., 2012). The alternative viewpoint favours variability
in tropical SST gradients, especially in the Pacific (Placzek et al., 2006, 2009). This group focuses on modern
analogue data, which show: (1) increased precipitation on
the Altiplano accompanies the strengthened trade-winds
of La Niña events; and (2) humidity in the Chaco lowlands plays a role as a moisture source, particularly for the
southern Altiplano (Quade et al., 2008).
The radiocarbon-dated stratigraphic framework presented herein allows the PB to be placed in a regional palaeoenvironmental context. According to our current age
model, the deepest palaeolake existed in the PB from ca.
44–37 cal ka BP, roughly coeval with the appearance of
the Minchin palaeolake in the Uyuni and Titicaca Basins
(Baker et al., 2001a, b; Chepstow-Lusty et al., 2005).
Similarly, water levels in the Salar de Hombre Muerto
and the Rıo Desaguadero Valley are inferred to have been
higher during this time (Godfrey et al., 2003; Rigsby
et al., 2005). This time interval spans Heinrich event 4,
which is very closely correlated with high precipitation
and a strong SASM observed in speleothem records from
Peru (Kanner et al., 2012).
A saline palaeolake is interpreted to have occupied the
PB from ca. 26 to 19 cal ka BP (Fig. 7). During this time,
the Sajsi palaeolake occupied the Uyuni basin, and water
levels were higher in the Salar de Atacama (Bobst et al.,
2001; Blard et al., 2011). Evidence of large lakes that correlate in time with the Tauca and Coipasa phases is conspicuously absent from the PB. Instead, these intervals,
and indeed all of the terminal Pleistocene and early Holocene, are overprinted by evidence of basin floor desiccation
and reworking associated with the dry mudflat (Facies C)
at the top of Unit II. Modification of the PB by wind is
clear in Facies C, and the OM geochemistry of these sediments is consistent with oxidation of previously deposited
lake sediments (Fig. 3b). Several intervals of aridity have
been inferred for the Altiplano/Puna, but most prominent
and relevant to the PB seems to be associated with an insolation minimum from ca. 8.0 to 4.0 cal ka BP (Mayle &
Power, 2008). The impacts of reduced insolation and a
northerly Intertropical Convergence Zone position on
monsoon precipitation were widely felt in tropical South
America, and expressed in proxy records from Amazonia,
the Pantanal, the Altiplano and mountain glaciers (Mayle
& Power, 2008; Burns, 2011; McGlue et al., 2012b).
Two modes of tectonic deformation are important for
lacustrine deposition in the PB. First, the formation of
the PB by thrust faulting and flexural subsidence controls
general aspects of basin geometry and the initial amount
of accommodation space. The parallel bounding faults
that make up the flanks of the PB constrain the size, depth
and hydrology of potential lake systems. The spacing
between the west-verging thrusts of the PB (<30 km)
indicates that palaeolakes most likely exhibited a low
width-to-depth ratio. These limnological conditions may
have helped to promote water column stratification and
preservation of both fine laminations and OM during
lake-level highstands (Katz, 1995). For the PB, the NNE
trend of boundary thrusts also limits the influence of easterly winds on the basin floor, reducing the fetch of any
large palaeolake. Facies data suggest that the deepest palaeolake in the PB (ca. 43–37 cal ka BP; Fig. 7) may have
been seasonally stratified, but placing constraints on the
surface elevation of this lake remains a challenge. If recent
tectonic movements have not altered the basin margins,
then ca. 35 m of accommodation space is available to hold
a hydrologically closed lake in the PB, given its present
day morphology (Fig. 1b). When hydroclimatic conditions allow this depth threshold to be crossed, spillover to
the northeast will occur and an open hydrologic system
will form. We hypothesize that the observed asymmetric
stratigraphy dominated by regressive deposits resulted
from lake levels remaining below the spill point elevation.
A prominent shoreline would be expected at a level near
the PB spill elevation, if palaeolakes had been persistently
open (e.g. Placzek et al., 2009). Although lower elevation
palaeo-shorelines exist in the basin, such features have
not been discovered near ca. 3695 m a.s.l.
In the broadest sense, the Andes form a barrier to eastern moisture sources, which has important implications
for the sediment plus water availability for all basins on
the orogen. Very few natural lakes exist in the wedgetop
basins and valleys of the Andean foothills (elevations
1500 m), where precipitation is relatively high and
dense Yungas evergreen forests are the norm. This is
because excess sediment plus water usually overwhelms
available accommodation space in these settings, leading
to at most the development of small ( 1 km2), shallow,
freshwater ponds or diffuse wetlands that accumulate thin
sapropels and clays. By contrast, large lake development
is possible in basins high atop the arid orogen, due to the
preservation of accommodation space from sediment
© 2013 The Authors
Basin Research © 2013 John Wiley & Sons Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists
M.M. McGlue et al.
starvation. Similar dynamics have also been observed in
Neogene Andean strata, which document the expansion
of palaeolakes when climatic (aridity) and tectonic (episodic thrusting) conditions reduced sediment supply to
the Bermejo Basin (Jordan et al., 2001).
Gangui (1998) noted that formation of the PB occurred
during the Oligocene, and the mid-Miocene (ca. 14 Ma)
carbonate-evaporite sediments of the Cara Cara Formation described by Cladouhos et al. (1994) attest to the
presence of another underfilled lake in the basin early in
its history. The geography of these outcrops, the geography of these outcrops suggests that this Miocene palaeolake may have been larger and deeper than the Quaternary
lakes identified in our core records. Seismic reflection data
suggest that sedimentation in the Miocene was syn-orogenic (Gangui, 1998), and thus the development of
accommodation space, coupled with negative P-E, most
likely led to basin underfilling and large lake development.
The largest palaeolake in the PB (ca. 43–37 cal ka BP;
Fig. 7) shares a number of similarities with late Oligocene-early Miocene lake deposits mapped by Horton
(1998). In Bolivia, wedgetop basins formed by westverging thrusts also evolved beginning in the Oligocene,
and mudstone-dominated lake beds with subordinate
carbonates and evaporites have been identified in some
of these systems (Horton, 1998). These small lakes
formed when fault-related surface deformation was at a
minimum, and lake expansion most likely followed
reduced sediment supply and relatively dry climatic
The second type of tectonic control important for
lake evolution in the PB is extension, as deformation
related to normal faulting could have impacted the
amount of accommodation space available for lakes or
influenced hydrologic networks. Pliocene-Quaternary
normal faults with variable offsets (up to ca. 5 m) have
been documented on the western side of the PB (Cladouhos et al., 1994). Aeromagnetic data have also
revealed basin-crossing transtensional lineaments
(Fig. 1a) that have been linked to reactivation of thrust
structures during Miocene volcanism on the PB’s
southern margin (Caffe et al., 2002). These extensional
structures could have influenced the character of Quaternary lakes in the PB through minor increases in
basin floor gradient and accommodation space. Another
possibility is that extension is responsible for altering
groundwater flow paths, which could help explain the
stark differences between Miocene and Quaternary lake
sediments in the PB. Notably, groundwater discharge
features in the palaeolake strata of the past 43 cal kyr
are absent, whereas the bedded evaporites of the Cara
Cara Formation suggest that lake hydrology was supported by groundwater in the Miocene. Horton (2012)
noted that structural style and depositional patterns in
a piggyback basin may evolve as the orogen advances
and the basin is incorporated into the hinterland
region. This process could affect many Puna basins,
due to extension in the thrust-belt hinterland proposed
to follow eclogite root foundering in Cordilleran orogenic systems (DeCelles et al., 2009).
(1) Five lithostratigraphic units (V–I; lake-plain/littoral,
profundal, palustrine, saline lake, and playa facies
associations respectively) identified in radiocarbondated sediment cores record sedimentation in the PB
over the past ca. 43 cal kyr. Pozuelos Basin facies are
characterized by an asymmetric vertical stacking pattern, where thick regressive facies overlie relatively
thin transgressive and highstand facies. The stratigraphy reflects the expansion and contraction of palaeolakes in an underfilled basin. Long-term (104 years)
sedimentation rates are relatively consistent among
the PB cores, whereas short-term (102–104 years)
rates are highly variable, which imply an incomplete
stratigraphic record that is compatible with facies
(2) Average TOC concentrations are highest in the profundal facies association (Unit IV), which suggests
that the production and preservation of OM were
greatest during maximum lake expansion and P-E.
Type II (mixed algal and macrophyte OM) kerogen
prevails throughout the late Quaternary PB record,
most likely a consequence of low accommodation
space and shallow bathymetry that characterized each
of the palaeolake environments. Siliciclastic dilution
was minimal, but organic facies development may
have been impacted by oxidation during prolonged
regressive phases.
(3) Late Quaternary lake dynamics and PB underfilling
were controlled by both tectonics and climate. Climate, through its effect on P-E and sediment supply, influenced the development of lakes of varying
character and water chemistry, some of which correlate in time with well-known Pleistocene palaeolakes in other basins on the Altiplano. The
structural configuration of the PB constrains the
morphology and fetch of highstand lakes, potentially promoting water column stratification and
development of laminated, organic-rich sediments
during highstands. Accommodation space maintained from piggyback basin-forming flexural subsidence and new space created as the PB evolved in
the Andean hinterland (principally through normal
faulting) likewise impacted lake hydrology and
This research was supported by the National Science
Foundation (Award 0542993), American Chemical Society (PRF grant 45910-AC8), ExxonMobil and small
© 2013 The Authors
Basin Research © 2013 John Wiley & Sons Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists
Puna basin lacustrine deposystems
grants from Sigma Xi, GSA, and AAPG to the first
author. L. Lupo and R.G. Cortes of the Universidad Nacional de Jujuy provided coring equipment. E. Piovano
and A. Kirschbaum arranged logistics and permitting. We
heartily thank J. Omarini, C. Gans, E. Gleason, M. Barrionuevo, M. Ayendez, D. Mu~
noz, and the staff of PN
Laguna de los Pozuelos for their assistance in the field.
The staff of LacCore graciously assisted with all aspects
of core processing and archiving. J. Wood drafted the palaeogeographical maps. Special thanks to our reviewers for
their thoughtful comments. Any use of trade, product, or
firm names is for descriptive purposes only and does not
imply endorsement by the U.S. Government.
Additional Supporting Information may be found in the
online version of this article:
Figure S1. Crossplot of sedimentation rate
(mm year 1) versus log time (years). Recent sedimentation rate data derived from radioisotopes are from
McGlue et al. (2012a). Long-term sedimentation rates
were calculated using basal 14C ages and a constant rate of
accumulation over the length of the dated interval (i.e.
not including coarse-grained basal units that lack direct
age control). Compaction is assumed to be negligible.
Average lacustrine sedimentation rate adapted from Cohen (2003). Late Quaternary rates suggest a punctuated
stratal record, which is compatible with facies observations in Unit II.
Figure S2. Shuttle Radar Topography Mission digital
elevation model of tropical and sub tropical South America, illustrating the locations of several palaeolake basins
mentioned in the text. 1, Salar de Hombre Muerto,
Argentina. 2, Salar de Atacama, Chile. 3, Pozuelos Basin,
Argentina. 4, Oligocene wedge-top basins, Bolivia (Horton, 1998). 5, Salar de Uyuni, Bolivia. 6, Rio Desaguadero
valley, Bolivia. 7, Lake Titicaca, Peru/Bolivia.
Table S1. Summary of core locations, lengths, and age
ranges used in this study.
Table S2. Radiocarbon geochronology of cores collected from the Pozuelos Basin.
Table S3. Lithofacies encountered in Cores 2A, 3A,
4A and 6A. Note that elemental data for Facies B (marked
by *) includes values previously published in McGlue
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Manuscript received 09 August 2012; In revised form 21
February 2013; Manuscript accepted 22 February 2013.
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