CLIMATE AND TECTONIC CONTROLS ON SEDIMENTATION AND DEFORMATION IN THE FIAMBALÁ BASIN OF THE SOUTHERN PUNA PLATEAU, NORTHWEST ARGENTINA
A Thesis
Presented in Partial Fulfillment of the Requirements for
the Degree of Master of Science in the
Graduate School of The Ohio State University
By
Heather Meybel McPherson, B.S.
* * * * *
The Ohio State University 2008
Master’s Examination Committee:
Dr. Lindsay Schoenbohm, Advisor
Dr. Lawrence Krissek Approved By:
Dr. Terry Wilson ______Advisor Geological Sciences Graduate Program
Copyright by Heather Meybel McPherson 2008
ABSTRACT
Here, we elucidate the relationship between tectonics and climate and their influence on sedimentation, erosion, deformation, coarsening of clastic sediment, regional uplift, and aridification by investigating Punaschotter Conglomerates of the Fiambalá Basin in
Northwest Argentina through mapping, cross-section construction and U-Pb dating of interbedded ashes. U-Pb zircon dating of seven interbedded ashes and ignimbrites from within the Punaschotter Conglomerates reveal deposition beginning at 4.02 ± 0.04 Ma.
Comparisons to regional basins show that deposition in the Toro and El Cajon Basins
preceded deposition in the Fiambalá Basin, while deposition in the Santa Maria,
Angastaco, and Hualfín-Río Basins is significantly younger. Regional conglomerate
deposition is asynchronous between basins, and with both global and regional climate
change. Therefore, we find that deposition is most likely the result of increased
aridification driven by regional uplift of basin-bounding ranges, and enhanced by
regional and global climate cooling between 6.4-2.6 Ma and 2-4 Ma, respectively.
ii ACKNOWLEDGMENTS
My sincere gratitude to Lindsay Schoenbohm, my advisor, for intellectual support,
guidance, and encouragement during my research and completion of this thesis.
Many thanks to Terry Wilson and Larry Krissek, my committee members, for
intellectual discussions, advice, and assistance in completing this thesis.
With much appreciation and many thanks to Juan Palevecino for his friendship, humor, guidance, and assistance in the field and around Argentina.
Thank you to Jonathan Pratt for assistance with sample preparation and analysis at
The Ohio State University and UCLA, and to Fritz Hubacher and Jeff Linder for
guidance in the mineral separation lab at The Ohio State University.
A special thank you to Ben Kirby and Wendy Bohon for their friendship, intellect, and humor over the past two years.
With utmost appreciation for my friends and family for their support during this very important personal and intellectual endeavour.
iii VITA
June 22, 1979……………………………………………Born - San Salvador, El Salvador
2001…………………………………………………….B.S. Geology, Denison University
2002-2006……………………………….……..Hydrogeologist, Eagon & Associates, Inc.
2006-present…………………………..……...Graduate Research and Teaching Associate The Ohio State University
PUBLICATIONS
1. McPherson, H., Schoenbohm, L., Carrapa, B., Schmidt, A., and Bohon, W., (2007), Climate and Tectonics Controlling Sedimentation in the Fiambala Basin, Northwest Argentina, Eos Trans. AGU, 88 (52), Fall Meet. Supple., Abstract T23D-1655.
2. Schoenbohm, L., Mortimer, E., Strecker, M., McPherson, H., Pratt, J., (2007), Deposition and Deformation in the El Cajon Basin, NW Argentina: a record of climate change, plateau growth and foreland fragmentation, Eos Trans. AGU, 88 (52), Fall Meet. Supple., Abstract T12D-06.
3. Bohon, W., Schoenbohm, L., Brooks, B., Costa, C., and McPherson, H., (2007), Drainage Analysis and Fluvial Terrace Reconstruction: Assessing Blind Thrust Hazards, Montecitos Anticline, Mendoza, Argentina, Eos Trans. AGU, 88 (52), Fall Meet. Supple., Abstract T23D-1652.
4. McPherson, H., and Hawkins, D., 2002, The June 1886 Eruption of Mount Tarawera, North Island, New Zealand, Denison Journal of the Geosciences, vol. 16, p. 24.
FIELDS OF STUDY
Major Field: Geological Sciences
iv TABLE OF CONTENTS
Page Abstract…………………………………………………………………………………....ii Acknowledgments…………………………………………………………………..……iii Vita………………………………………………………………………………………..iv List of Tables…………………………………………………………………………….vii List of Figures…………………………………………………………………………...viii
Chapters:
1. Background………………………………………………………………………..1 1.1 Introduction……………………………………………………………………1 1.2 Motivation……………………………………………………………………..2 1.3 Tectonics Driving Climate Change……………………………………………3 1.4 Climate Change Driving Increases in Sedimentation and Erosion...………….6
2. Climate and Tectonic Comparison within the Fiambalá Basin, Northwest Argentina………………………………………………………………………...... 9 2.1 Introduction……………………………………………………………………9 2.2 Regional Geologic Setting…………………………………………………...12 2.2.1 Puna Plateau……………………………………………………………12 2.2.2 Intramontane Basins……………………………………………………14 2.2.3 Fiambalá Basin…………………………………………………………16 2.2.4 Other Regional Basins…………………………………………………24
2.3 Climatic Setting……………………………………………………………...23 2.3.1 Global Cenozoic Climate Change……………………………………...23 2.3.2 Orbital Forcing of Climate……………………………………………..30 2.3.3 Continental Scale Tectonic Forcing of Climate………………………..32 2.3.4 Local Tectonic Forcing of Climate…………………………………….36
2.4 Methods………………………………………………………………………39
2.5 Results………………………………………………………………………..44
2.6 Discussion……………………………………………………………………54 2.6.1 Age of the Punaschotter Conglomerates in the Fiambalá Basin……..54 2.6.2 Deposition, Folding, and Faulting in the Fiambalá Basin…………...61
v 2.6.3 Geochronology of the Punaschotter Conglomerates Along the Puna Plateau………………………………………………………….…………….64 2.6.4 Comparison to Regional Basins…………………………….………..68 2.6.5 Migration of a Tectonic Orogen………..…………………………....70
2.7 Comparison to Climate and Tectonics……………………………….………71 2.7.1 The Fiambalá Basin Punaschotter Conglomerates…………….………71 2.7.2 Regional Basin Conglomerates……………………………………..….72
2.8 Synthesis……………………………………………………………………..75
3. Conclusion…………………………………………………………………………….78
3.1 Summary………………………………………………………………….….78 3.2 Limitations……………………………………………………………….…..80 3.3 Data Issues……………………………………………………………….…..81 3.4 Future Work………………………………………………………………….82
References Cited…………………………………………………………………………83
vi LIST OF TABLES
Table Page
1. Climate Summary – Past 10 Ma……………………………………………..…..29
2. U-Pb Geochronology Results……………………………………………..……..42
vii LIST OF FIGURES
Figure Page
1. The Altiplano-Puna Plateau of South America…………………….…………….13
2. Regional basins and associated thrust and reverse faults along the margin of the Puna Plateau………………………………………………………………….…..15
3. The Fiambalá Basin……………………………………….……….………….…17
4. Tamberia Formation and Guanchin Formation………………..……………..….20
5. The Punaschotter Conglomerates …………………………………………….…21
6. Fiambalá Stratigraphy……………………………………………………………23
7. Age of Regional Exhumation, Uplift and Deformation …….……………..……25
8. Cross sections locations within the Fiambalá Basin…………………………..…40
9. Map and cross section A-A’...………………………………………..………….45
10. Map and cross section B-B’……………………………….……………………..47
11. Map and cross section C-C’……………………………………………………...48
12. Map and cross section E-E’……………………………………………………...50
13. Map and cross section D-D’……………………………………………………. 51
14. Map and cross section F-F’……………………………………………………....52
15. Map and cross section from Carrapa et al. (in press)…...…………….………….53
16. Punaschotter-Guanchin contact………………………………..……………...…56
17. Paleowatersheds in the Fiambalá Basin…………………………..………...……60
viii 18. Geologic and structural mapping of the Fiambalá Basin………………………...62
19. Age of the regional Punaschotter Conglomerates………………………………..65
ix
CHAPTER 1
BACKGROUND
1.1 INTRODUCTION
Tectonics and climate can each independently affect the landscape by changing rates
of erosion and sedimentation, through uplift and subsequent shedding of sediment from
nearby mountain ranges, by changing depositional patterns, and by causing aridification.
However, tectonics and climate cause similar changes in the landscape, making it
difficult to distinguish which of the two is the principal driving force in a particular case.
This study seeks to elucidate the relationship between tectonics and climate and their
influence on sedimentation, erosion, and deformation during the late Miocene-Pliocene in
the Fiambalá Basin of Northwest Argentina. This is done by structural, geologic, and
geomorphic mapping, and U-Pb zircon dating of interbedded ashes within the
Punaschotter Conglomerates shed from the nearby Puna Plateau. Our results are
compared with those of other basins in the region to more accurately evaluate and
constrain timing of tectonic and climatic influences.
This paper is composed of three chapters. The opening chapter provides a background
to the problem of distinguishing the interlayered, interdependent effects of climate and tectonics on erosion, exhumation and the sedimentary record. The second chapter presents the details found by this study for the timing of Punaschotter deposition in the
1 Fiambalá Basin, makes a comparison to other basins in the region, and discusses the overall implications for both tectonics and climate change related to this study. The final chapter is a conclusion which summarizes the findings of this study, discusses the limitations of our research and problems with the data, and makes suggestions for future study in the Fiambalá Basin.
1.2 MOTIVATION
Increased sedimentation and erosion rates, and increased clastic grain size in the last
2-4 Ma (Peizhen et al., 2001) have commonly been ascribed to late Cenozoic uplift of mountain ranges worldwide. In 1990, Molnar and England proposed an alternative to this hypothesis and suggested that such increases could be the result of global climate cooling by means of increased glacial erosion, increased erosion of continental shelves due to lowered sea levels, and climate oscillations driven by variations in orbital parameters
(e.g. Molnar, 2004; Peizhen et al., 2001). Evidence used to infer late Cenozoic uplift was, in fact, evidence of late Cenozoic climate change. Molnar and England (1990) suggested that the interpretation of geomorphic, sedimentological and paleobotanical evidence, used to infer late Cenozoic uplift, is exaggerated and overlooks evidence of global cooling. A geomorphic process commonly used to infer recent uplift is the sharp incision by streams and rivers into surfaces, resulting in increased erosion rates.
However, climate change can also be a significant contributor to accelerated erosion rates through glacial erosion (Molnar and England, 1990). Thick deposits of late Cenozoic conglomerates close to steep mountain ranges are often attributed to tectonic activity, but abrupt coarsening of sediment could also be the result of infrequent but large storms due
2 to increased global aridity (Frostick and Reid, 1989). Many paleobotanical inferences of recent uplift ignore global cooling altogether, and may be highly subjective if paleoelevations and paleoclimates are determined from present-day fossil assemblages and elevations (Molnar and England, 1990). Molnar and England also contend that the term “uplift” is used inconsistently and is often confused; uplift of the Earth’s surface due to isostatic rebound and uplift due to exhumation are different tectonic components. They acknowledge, however, that it is difficult to distinguish a local climate change due to global processes from one due to local uplift.
This “chicken and egg” conundrum, as it has come to be known, poses the question:
Has tectonic uplift during the late Cenozoic driven increases in sedimentation and erosion rates and clastic grain size, and driven climate change, or has late Cenozoic climate change driven these increases, and perhaps caused tectonic uplift? Are increases in sedimentation and erosion rates, as well as clastic grain size, the result of regional tectonic forces or regional or global climatic effects? Further complicating the issue is the fact that global climate change has distinct regional expressions and may not have the same effects worldwide (e.g. Mix, 1995). Studies in support of each argument are equally convincing, yet caveats remain. No argument can completely rule out the impact of either climate or tectonics, making it more difficult to distinguish the effects of each on the landscape.
1.3 TECTONICS DRIVING CLIMATE CHANGE
Climate models employed by Ruddiman, Raymo, and others make a strong case for tectonic uplift driving climate change. Raymo and Ruddiman (1992) explore late
3 Cenozoic uplift of the Tibetan Plateau as the driving force for global cooling and ice
sheet expansion. The great height and width of the Tibetan Plateau impacted monsoonal
patterns of the region and influenced atmospheric circulation throughout the Northern
Hemisphere (e.g. Ruddiman and Raymo, 1988; Ruddiman and Kutzbach, 1989;
Ruddiman and Prell, 1997). Increased elevation, particularly widespread regional uplift,
favors increased snow accumulation, extending the length of winter and increasing the
albedo, working in an albedo-temperature feedback (Ruddiman and Prell, 1997), and
creating a positive feedback between tectonics and climatic cooling (e.g. Birchfield and
Weeterman, 1983; Raymo and Ruddiman, 1992). General circulation models (GCM) run
by Ruddiman and Kutzbach (1989) show that, while variable landmass topography does
affect precipitation, temperature, and paleobotany, high topography alone will not drive
the expansion of ice sheets in both hemispheres. Changes must also be present in
atmospheric circulation patterns, and positive feedback must occur between many
parameters. Molnar and England (1990) do not discount such models as evidence for
climate change over long periods of time. Changes in mean elevation, climate change and
the phenomena commonly used to infer uplift (geomorphic, sedimentological and
paleobotanical evidence) are intimately coupled to one another.
Additionally, physical erosion can cool global temperature by extracting CO2 from the atmosphere due to three tectonic mechanisms: 1) active faulting that exposes fresh unweathered rock, 2) high-altitude mechanical weathering due to steep slopes, lack of vegetation, and periglacial-glacial weathering, and 3) orographic rainfall on windward slopes of a mountain range. These three processes can cause increased rates of erosion at high elevations and provide abundant sediment, driving rapid chemical weathering at
4 lower elevations. The accelerated chemical weathering of silicate minerals, reactions that
absorb carbon dioxide, decrease the greenhouse effect and causes cooling (Raymo et al.,
1988).
In addition to the effect of plateau uplift on global climate, climate can be affected
locally by tectonic uplift, as has been well documented in the study region. Work carried
out along the eastern margin of the Puna Plateau by Kleinert and Strecker (2001), Sobel
and Strecker (2003), Carrapa et al. (2005, 2006), Coutand et al. (2006), and Deeken et al.
(2006) demonstrate that the uplift of the Puna Plateau and the eastward migration of the
Andean tectonic orogen created regional orographic barriers. The windward side of a
mountain range receives high amounts of precipitation while the leeward side receives
little, leading to increased aridity on the lee side. Apatite fission track thermochronology
has constrained exhumation and deformation of basin-bounding ranges adjacent to the
Puna Plateau to the late Eocene-mid Miocene, and continued uplift of these ranges to late
Miocene time, ~8-5 Ma (e.g. Carrapa et al., 2005, 2006; Deeken et al., 2006). Climate
change has been well documented on and around the Puna Plateau, with aridification being linked to the formation of internal drainage and deposition of halite and gypsum- bearing units between 24-15 Ma on the plateau (Alonso et al., 1991; Vandervoort et al.,
1995). Sediments in intramontane basins record climate change during Miocene-Pliocene time. Stark and Anzóegui (2001) and Coutand et al. (2006) document a change from aridity to wetter conditions at ~10 Ma and a transition from wetter conditions to semi- aridity at ~3.4-2.4 Ma based on depositional environments. Kleinert and Strecker (2001) use paleobotanical evidence to demonstrate that uplift of orographic barriers affected precipitation in and around the Santa Maria Basin throughout the last 12 Ma, with
5 significant aridification from ~3-2.5 Ma. Other studies find little evidence of wetter
conditions, but interpret increased aridity from ~10-3.7 Ma using sedimentological data
(Carrapa and Mortimer, in preparation; Carrapa et al., in press). In summary, evidence
from basins along the eastern margin of the Puna Plateau indicate that regional uplift of basin-bounding ranges was responsible for the creation of multiple orographic barriers, driving changes in precipitation and atmospheric circulation, ultimately leading to increased aridification by the Pliocene. Deposition of the Punaschotter Conglomerates may have been the result of increased local aridity driven by bounding range uplift along the margin of the Puna Plateau. In the following chapter, we discuss the mechanisms by which global cooling and aridity lead to increased erosion, sedimentation, and sediment grain size.
1.4 CLIMATE CHANGE DRIVING INCREASES IN SEDIMENTATION AND
EROSION
Strong evidence also exists in support of the proposal that global climate change
drove increases in sedimentation and erosion rates, and clastic grain size at 2-4 Ma.
Sediment accumulation increased dramatically both on land and in the oceans since at
least 4 Ma, and has been attributed to lowered sea levels along continental margins due to
increased continental glaciation (Hay et al., 1988). Peizhen et al. (2001) assert that the
only process capable of causing the global increase in sedimentation and erosion rates
and increased clastic grain size is global climate change. First, globally synchronous
movement of tectonic plates, possibly resulting in globally synchronous tectonic uplift,
has not occurred since ~9 Ma, as documented by astronomical calibrations of the
6 geomagnetic timescale (Krijgsman et al., 1999). Second, sedimentation rates increased in
both tectonically active and inactive regions (e.g. Molnar, 2004: Peizhen et al., 2001),
thus eliminating tectonic uplift as the sole driver of increases in sedimentation, erosion, and clastic grain size. Third, thick deposits of conglomerates often found close to steep mountain ranges could be the product of increased glacial erosion, and not tectonic activity. Most of the major mountain ranges thought to have risen in the late Cenozoic
were also glaciated; thus glaciation could also have been responsible for increased
sedimentation rates and clastic grain size (Molnar and England, 1990). Many of the
mountain ranges surrounding the Puna Plateau were once glaciated; deposition of the
Punaschotter Conglomerates may also have been the product of increased erosion and
sedimentation rates, due to increased glacial erosion.
It has been argued that climate drove changes in erosion and sedimentation rates in
the Andes, and may also have led to formation of the Puna Plateau (Lamb and Davis,
2003). Climate-induced sediment starvation on the western side of the Andes led to
higher shear stress and coupling in the subduction trench, causing more shortening in the
overriding plate, thus building the Andes. Lamb and Davis (2003) argue that the extreme
aridity of northern Chile and Peru is primarily the result of atmospheric and oceanic
circulation, deep ocean cooling, and the Central Andean rain shadow, driven by global
climate cooling. As the East and West Antarctic ice sheets expanded (~14 and ~6 Ma,
respectively), cooling the Pacific Ocean currents, the Andes were undergoing their most
extreme phase of shortening and rapid surface uplift (20-10 Ma) (Lamb and Davis, 2003).
Drying continued and hyperaridity developed in western South America by 3-4 Ma.
Studies by Hartley and Chong (2002), Hartley (2003), and Dunai et al. (2005) generally
7 support these climate changes based on sedimentological data indicating increased
aridity, and conclude that global climate change, not Andean uplift, was responsible for
aridification of western South America. In this scenario, deposition of the Punaschotter
Conglomerates would have been the result of Andean uplift and physical erosion. Here, however, the cause of deposition would be tectonic uplift driven indirectly by increased
aridity and climate change.
The following chapter presents the data and results found by this study and relates the
timing of Punaschotter deposition to both global and regional climate changes and to
regional tectonic uplift.
8
CHAPTER 2
CLIMATE AND TECTONIC COMPARISON WITHIN THE FIAMBALA BASIN, NORTHWEST ARGENTINA
2.1 INTRODUCTION
Global climate change during the Late Pliocene has been credited for a worldwide increase in erosion and sedimentation rates and a concurrent abrupt coarsening of clastic sediment (Molnar and England, 1990). Accumulation rates of terrestrial sediment increased by two to ten times worldwide at 2-4 Ma (Peizhen et al., 2001). These worldwide occurrences could be the result of three factors: 1) lowered sea-level exposing continental shelves, thus leading to higher rates of erosion, 2) increased erosion at high altitudes and latitudes due to glacial expansion, and 3) large climate oscillations operating on 20,000-40,000 year cycles which could have created a disequilibrium in the landscape, causing erosion rates to increase (Molnar, 2004). However, local structural deformation and regional tectonic uplift may also explain increased sedimentation rates in certain cases. Crustal deformation creates regionally elevated terrain that provides potential energy to rivers and glaciers, the main agents of erosion (Peizhen et al., 2001). Such differing interpretations have fueled an on-going debate about the relative roles of both climate and tectonics and degree to which each shape the landscape.
9 This study seeks to further elucidate the relationship between climate and tectonics by investigating the coarse clastic sediment shed from the Puna Plateau in northwestern
Argentina, known as the Punaschotter Conglomerates (Penck, 1920). Clastic sediment preserved within intramontane basins adjacent to plateaus offers a unique record of the timing and pattern of orogenic evolution and its relationship to tectonics and climate (e.g.
Hilley and Strecker, 2005; Carrapa et al., 2005). Thick deposits of upper Cenozoic conglomerate surround many steep mountain ranges, and the large cobbles that comprise the conglomerates imply transport along steep slopes (Molnar and England, 1990). While tectonic activity is often supposed as the mechanism causing abrupt clastic coarsening and increased conglomerate influx, climate change may also be responsible due to increased glaciations and stream potential (Molnar and England, 1990). These conglomerates may record global and/or regional climate change, tectonic activity, or both, and can serve as a proxy for tectonic activity worldwide.
Constraining the history of deposition and deformation within the Fiambalá Basin is vital to understanding the roles that both climate and tectonics play in basin evolution and sediment deposition. Clastic sediment preserved within the Fiambalá Basin may be a robust recorder of Miocene-Pliocene climate change, as well as the orogenic evolution of the Puna Plateau margin. By constraining the ages of bounding range deformation and
uplift, basin formation and drainage reorganization, and sediment deposition, we are able
to more accurately construct an evolutionary history for the active tectonic orogen. By
understanding the history and development of the Central Andean climatic regime for the
last 5-10 Ma, we can compare and contrast these climate changes to tectonic events.
Currently, the onset of Punaschotter Conglomerates deposition in the Fiambalá Basin has
10 been constrained between 3.77 ± 0.05 Ma and 5.23 ± 0.30 Ma (Carrapa et al., in press).
Deposition of these conglomerates thus require tighter constraint and further refinement,
which is the purpose of this study.
The age of the Punaschotter Conglomerates can be compared to conglomerate
deposition in other basins along the eastern Puna Plateau margin to determine if
deposition was synchronous with regional tectonic activity or global climate change. If
worldwide climate change between 2-4 Ma (Peizhen et al., 2001) was the trigger for
increases in clastic sediment accumulation, erosion rates, and clastic grain size, then the timing of regional exhumation and sediment deposition and erosion should correlate strongly with episodes of global climate change and be synchronous among basins along the eastern margin of the Puna Plateau. The global climate change discussed herein is often ascribed to the late Cenozoic, without definition of absolute dates. Here, we use 2-4
Ma, as referenced by Peizhen et al. (2001), as the time of late Cenozoic global climate change. If a climate and tectonic correlation is weak or non-existent, regional exhumation and sediment deposition will vary between individual basins, and deposition and erosion will be syn-tectonic.
The results of our study and our comparison to other regional basins lead us to
conclude that regional tectonics is the most probable driving force for increased erosion
and accumulation rates, and increased clastic grain size. Additionally, our study aids in
better understanding the tectonic and climatic development of the Puna Plateau and its
effect on the surrounding region.
11 2.2 REGIONAL GEOLOGIC SETTING
2.2.1 Puna Plateau
The Fiambalá Basin and other basins of interest in this study are located along the
southeastern-eastern margin of the Puna Plateau in northwestern Argentina. The Puna-
Altiplano Plateau is the second largest plateau on earth and extends for ~1800 km from
southern Peru to northern Argentina (Figure 1). It has a maximum width of 350-400 km
with an average elevation of ~ 4 km (Allmendinger et al., 1997). The Puna has an
average elevation nearly a kilometer higher than the Altiplano (Allmendinger et al.,
1997), attributed to greater thinning of the lithosphere beneath the Puna (Whitman et al.,
1996). The plateau formed by the subduction of the oceanic Nazca Plate beneath South
America. The subducting Nazca Plate changes dip along strike owing to plate subduction geometry and bathymetric relief being affected by the shape of the South American plate
(e.g. Cahill and Isacks, 1992). Beneath the Bolivian Altiplano (15°S - 22°S), the plate
dips at 30°, gradually transitioning underneath the Puna to a near-horizontal dip between
latitude 28°S - 32°S. Between 25-27° S there is a notable gap in the Wadati-Benioff
earthquake zone (Cahill and Isacks, 1992).
Subduction of the Nazca Plate caused extensive crustal shortening and thickening, lithospheric thinning and thermal softening, and magmatism (e.g. Allmendinger et al.,
1997). Regional deformation commenced in the Altiplano during the Paleocene-Eocene
(Lamb et al., 1997), while deformation dominated the Puna Plateau from the Eocene-
Oligocene until the Pliocene-Pleistocene time (Coutand et al., 2001), causing shortening
and uplift of basement blocks, particularly along the plateau margin and foreland regions
(e.g. Strecker et al., 1989; Coutand et al., 2006). The direction of contraction changed
12
Figure 1. Anaglyph, South America (PIA03389) showing the location and extent of the Puna-Altiplano Plateau in brackets. Figure from NASA/JPL/NIMA, February 2000.
13 from NW-SE to NE-SW at ~2-4 Ma, the result of plateau uplift and lateral extrusion of material from the plateau toward the south (Allmendinger, 1986). After this time, strike-
slip faulting and extensional faulting dominated plateau tectonics (Allmendinger et al.,
1997).
Internally, the plateau is broken into contractional “basins and ranges,” and ranges are
composed of deformed Paleozoic rocks (Allmendinger et al., 1997). The plateau is bounded to the southeast by high angle reverse faults of the Sierra Pampeanas basement
blocks, the Eastern Cordillera fold-and-thrust belt to the northeast, and a magmatic arc to the west. Mountain ranges in excess of 6000 m that were once sparsely glaciated surround the plateau to the north, east and west; from southwest to northeast lie high- elevation (3500-3600 m), internally drained intramontane basins. These basins are situated within the Sierras Pampeanas morphotectonic province (Strecker et al, 2007), the structural transition between the Puna mountain ranges, the Eastern Cordillera, and the
Santa Barbara thrust system (Strecker et al., 2006). The Sierras Pampeanas, located between 27° and 33° S latitude, are east of the amagmatic sector and are coincident with shallow subduction of the Nazca Plate (Jordan et al., 1983).
2.2.2 Intramontane Basins
Intramontane basins of the Puna Plateau margin include the Fiambalá, Hualfín-Río,
El Cajon-Campo del Arenal, Santa María, Angastaco, and Quebrada del Toro basins
(Figure 2a). These basins are characteristically arid and preserve thick Miocene to
Quaternary sedimentary strata which record timing of deposition and deformation, plateau exhumation and morphology, and basin compartmentalization, leading to
14
aco, and Quebrada e Puna Plateau in South America. a Maria, Angast a Maria, 06) and from Strecker et al. (2007). M) of the southeast margin of th of margin M) of the southeast fín, El Cajon-Campo del Arenal, Sant lt locations from Carrapa et al. (20 del Toro Basins shown. Figure 2b. Digital elevation map (SRTM) of thrust and reverse fault locations along thedel Toro Basins shown. Figure 2b. Digital elevation map (SRTM) Puna Plateau margin. Fau Locations of the Fiambalá, Rio-Hual a b Figure 2a. Digital elevation map (SRT
15 drainage reorganization. Basins are structurally controlled by reverse faulting, which
brings Paleozoic to Mesozoic basement rock over thick Miocene to Quaternary clastic
sequences (Figure 2b). Regional basins are structurally similar to the internally drained
basins within the plateau, but fluctuate spatially between internal drainage and open
drainage patterns with a connection to the foreland plateau (Strecker et al, 2006).
Consequently, basins record multiple sediment filling and evacuation events from Mio-
Pliocene to Quaternary time. Some of the basins may have been connected in a
continuous foreland basin in the past, but have since been compartmentalized by uplift of
intervening ranges (e.g. Marrett and Strecker, 2000; Hilley and Strecker, 2005). Though
timing varies from basin to basin, uplift generally commenced by the late Miocene (~8-5
Ma) after the last Neogene transgression into the Andean foreland and may have accelerated after ~4 Ma (e.g. Sobel and Strecker, 2003). After ~3 Ma, basins continued to be folded and partly overthrust by crystalline basement uplifts (Strecker et al., 1989).
2.2.3 Fiambalá Basin
This study focuses on the Fiambalá Basin, which is located at approximately 27° 45’
S and 67° 45’ W with a basin floor elevation of ~1650 m (Carrapa et al., 2006). The basin is bounded to the west by east-vergent reverse faults (Figure 3) which place rocks of the
Sierra de las Planchadas of the Famatina Range on top of Miocene-Pliocene sedimentary strata. One major reverse fault runs the length of the basin, though numerous smaller reverse faults and blind faults also accommodate displacement (Carrapa et al., in press).
The Fiambalá Basin is bounded to the east by the Sierra de Fiambalá range. Rather than reverse-thrust over strata of the Fiambalá basin, the Sierra de Fiambalá range is back-
16
Figure 3. Digital Elevation map (SRTM) of the Fiambalá Basin and basin- bounding thrust and reverse faults. Fault locations from Carrapa et al. (2006).
17 tilted by faults further east. This range abuts the Puna Plateau to the north (Cerro Negro)
and is open to the south.
The Sierra de las Planchadas to the west of the Fiambalá Basin are composed of
Precambrian granite, aplite and diorite; Ordovician-Devonian siltstones and mudstones;
Permian volcanics, clastic sedimentary rocks and basaltic conglomerates; and Tertiary andesitic volcanics (Turner, 1967). To the north-northeast, the southeastern segment of the Puna Plateau is composed of Cambrian migmatite, schist, granite, and orthogneiss
(Turner, 1967).
Thermochronologic data show that the western bounding ranges – the Sierras de las
Planchadas - underwent deformation during the Eocene and exhumation during the early- middle Miocene (Carrapa et al., 2006), while deformation and exhumation of the eastern
Sierra de Fiambalá and northern Cerro Negro ranges began during the Oligocene-
Miocene (Coutand et al., 2001; Carrapa et al., 2006). Inferred uplift commenced in the late Miocene (~7 Ma). Deformation style is consistent through the basin, with fault and fold axis orientations documenting an older phase of NW-SE contraction followed by a younger phase of NE-SW contraction (Carrapa et al., in press). A NE-striking, NW- dipping normal fault indicates extension in the northern end of the basin during Plio-
Quaternary time (Schoenbohm et al., 2005). This kinematic shift could have been the result of a change in plate convergence direction (Marrett and Strecker, 2000), delamination of the mantle lithosphere beneath the plateau, N-S extensional collapse of the plateau, lower crustal flow, or a combination of all (Schoenbohm et al., 2005).
Pre-Miocene sedimentary rocks in the Fiambalá region suggest an older basin history
than that presently preserved by sedimentary sequences (e.g. Carrapa et al., in press;
18 Turner, 1967). Apatite fission track dating from basin-bounding ranges at 28° S (Sierra
La Maz, Sierra Umango, and Sierra de Famatina) document that these ranges were
covered by an ~11 km thick pile of upper Paleozoic to Neogene sediment at ~ 20 Ma
(Coutand et al., 2006; Coughlin et al., 1998). Thermal modeling of AFT detrital
populations confirms this by showing that key sources for basin sediment were covered
by thick sedimentary sequences until ~20 Ma (Carrapa et al, 2006). The Fiambalá Basin
may have once been connected to the Corral Quemado (Hualfín-Río) Basin. During the
late Miocene, the basin was disrupted twice by eastward fault propagation and
deformation (~9 and 6 Ma), which occurred simultaneously with paleodrainage
reorganization, basin compartmentalization, and relief development in the southern Puna
margin (Carrapa et al., 2006). Thus, the sedimentary record of the present Fiambalá Basin
is likely only a partial representation of its entire sedimentary history.
The Fiambalá Basin contains over 4 km of upper Miocene-Pliocene deposits (Turner,
1967). The Tamberia, the Guanchin, and the Punaschotter Formations record exhumation
and erosion of the surrounding ranges (e.g. Tabutt, 1986; Jordan and Alonso, 1987;
Carrapa et al., 2006). The Tamberia Formation (ca. 2000m) (Figure 4) consists of red sandstones, siltstones, and clast-supported conglomerates, and has an overall tabular and
coarse character which coarsens upward (Carrapa et al., in press). The Guanchin
Formation (ca. 1500 m) (Figure 4) is composed of mudstones, siltstones, medium to coarse-grained sandstones, and intercalated with fine and medium-grained
conglomerates; overall, the unit coarsens and thickens upwards (Carrapa et al., in press).
The Punaschotter Formation (ca. 500 m) (Figure 5) is composed of medium to coarse- grained conglomerates which are clast-supported and poorly sorted (Carrapa et al., in
19
Figure 4. Tamberia Formation (top) and Guanchin Formation (bottom).
20
Figure 5. Examples of the Punaschotter Conglomerates throughout the Fiambalá Basin. In the northern section of the basin, conglomerates form high terraces (top left). Sedimentation of conglomerates ranges from well sorted (bottom left) to poorly sorted (top and bottom right).
21 press). Collectively, these three units comprise an overall upward-coarsening succession
of ephemeral braided fluvial and alluvial fan deposits reflecting progradation and syn-
tectonic deformation (Carrapa et al., 2006; Hauer et al., 2007). A simplified stratigraphic
column of these three units by Carrapa et al. (in press) is shown in Figure 6.
The Guanchin and Punaschotter Formations crop out throughout the basin, while the
Tamberia Formation crops out in the southern portion of the basin. To the south, these
late Miocene-Pliocene units are highly faulted and folded and, in places, overturned. The
Punaschotter Formation unconformably overlies the Guanchin Formation in some places, and is transitional with the Guanchin Formation in other locations. The conglomerates contain numerous interbedded ash layers used for zircon U-Pb geochronologic dating.
U-Pb dating by Carrapa et al. (in press) of three ashes interbedded in the Punaschotter
Formation along the Rio Guanchin transect gave ages of 3.69 ± 0.06 Ma, 3.74 ± 0.05 Ma,
and 3.77 ± 0.05 Ma, from top to bottom (see figure 15a for sample locations). An
additional ash in the Punaschotter Conglomerates along the Northern transect (see figure
13a for location), yielded an age of 3.05 ± 0.44 Ma (Carrapa et al., in press). Two ashes
from within the Guanchin Formation were collected along the Rio Guanchin transect
(Figure 15a), yielding ages of 5.51 ± 0.15 Ma and 5.23 ± 0.30 Ma (Carrapa et al., in press). One ash was collected in the Guanchin Formation along the Northern transect
(Figure 13a), recording an age of 5.90 ± 0.20 Ma (Carrapa et al., in press). Based on youngest grain analysis for one of the samples, an age bracket of 3.77 ± 0.05 Ma to 5.23
± 0.30 Ma was determined for onset of deposition of the Punaschotter Conglomerates
(Carrapa et al., in press).
22
Figure 6. Simplified stratigraphic column by Carrapa et al. (in press), from the Rio Guanchin transect. ZFT (zircon fission track) ages from Tabutt et al. (1986). pmg = paleomagnetic data from Reynolds (1987). (U/Pb)* ages are from dated ashes by Carrapa et al. (in press). The best age for the Guanchin Formation (029) and one single mean age for the Punaschotter Formation are reported. mrl = marlstone, fs=fine sandstone, Gr = granule, cgl=conglomerate. Sedimentation rates are calculated using the ca. 9 Ma constraint from paleomagnetostratigraphy (Reynolds, 1987) for the base of the Tamberia Formation and new U-Pb ages for the rest of the stratigraphy.
23 The Punaschotter Conglomerates are interpreted by Carrapa et al. (in press) as debris
flow deposits in a proximal alluvial fan environment that received sediment from a
variety of sources in the surrounding ranges. Paleocurrent data and facies analysis
indicate source regions first from the north and east along the Guanchin transect, due to
erosion off the Puna margin, and later from sources to the west, north, and east for
Punaschotter deposition further east in the Fiambalá Basin due to expansion of the
drainage system (Carrapa et al., in press). The paleodrainage system reorganized and
expanded throughout the basin at ~6 Ma. This was ultimately the result of eastward
foreland propagation, which caused eastward migration of deformation and growth of the
Puna Plateau to the north (Carrapa et al., in press). The partially closed basin may have
overfilled by 3.7 Ma, preventing the establishment of internal drainage (Carrapa et al., in
press).
2.2.4 Other Regional Basins
Basin adjacent to the Puna Plateau have well-documented age constraints for
deformation, exhumation, and uplift of bounding ranges (Figure 7). The Hualfín-Río, El
Cajon-Campo del Arenal, Santa María, Angastaco, and Quebrada del Toro basins are
regional intramontane basins that border the Puna Plateau to the south and east. Uplift of
local mountain ranges that border each basin, hereafter known as basin-bounding ranges,
is likely to have caused the deposition of the Punaschotter Conglomerates; by
constraining the uplift history of these ranges, we gain a better understanding of tectonic development in and around the Puna Plateau and can more accurately assess the possible cause of Punaschotter deposition.
24
ferred uplift; on and inferred uplift, Sierra de rra Santa Rosa de Tastil; AFT; ra Pasha; inferred from transition Ma), Sierras de las Planchadas Ma), Sierras humation, northern Sierra Quilmes humation, ift, Cerro Negro and Alto Grande; ) deformation and exhumation, ents, Hilley and Strecker, 2005. cene (5.4-2.4 Ma) in tion, Sierra Aconquija; AFT; Sobel ) deformation and exhumation, exhumation, Oligocene-Miocene and uplift, Cumbres Luracatao Cumbres Luracatao uplift, and nd et al., 2006.; Deeken 2006. al., 2003; Mortimer et 2006. rred deformation and uplift, rtimer et al., 2006. rrapa et al., 2005. 1. Eocene-Oligocene (~44-35 Ma and Filo Negro; AFT; Carrapa et al., 2005; 2006. 2. Oligocene-Miocene (~23-15 Ma inferred upl (>6 Ma) Late Miocene 3. Late Miocene (~6 Ma) exhuma and Strecker, 2003; Mo 4. Late Eocene-Early Oligocene (38-29 Ma) exhumation, Sierra de Chango Real; AFT; Coutand et al., 2001. 5. Late Miocene (7-6 Ma) infe et Sierra Quilmes; AFT; Sobel 6. Late Miocene (9-5 Ma) exhumation, Cumbres Calchaquies; AFT; and Strecker, 2003. Mortimer et al., 2006. ; Sobel Ma) ex 7. Mid-Late Miocene (12-7 range; AFT; Butz et al., 1995; Deeken 2006. 8. Late Oligocene (29-24 Ma) exhumati Calalaste; AFT; Ca 9. Oligocene (30 Ma) deformation, Sie Andriessen and Reutter, 1994. 10. Late Eocene-early Oligocene (25-15 Ma), inferred deformation (Eastern Cordillera); AFT; Couta Sierra de los Colorados 11. Late Miocene (9 Ma) deformation, (Santa Barbara ranges), Early Plio AFT; Coutand et al., 2006. 12. Late Miocene (8-6 Ma) uplift Sier of sediment depositional environm inferred uplift by Late Miocene (~7 AFT; Carrapa et al., 2006. in. g lateau mar lateau Figure 7. Age of regional exhumation, inferred uplift and deformation of the Puna Plateau basin-bounding ranges. Colored boxes (yellow) (red) to younger older illustrate deformation, exhumation and uplift across the p
25 The Hualfín-Río Basin is located approximately 75 km northeast of the Fiambalá
Basin and lies to the west-southwest of the El Cajon-Campo del Arenal basins. The basin
is divided by the Sierra de Hualfín Anticline, estimated to have uplifted between 10-2 Ma
(Allmendinger, 1986; Garcia and Davis, 2004). Both portions of the Hualfín-Río Basin
are bounded by west-dipping high-angle reverse faults. To its northeast lies the Chango
Real range, exhumed in the late Eocene-early Oligocene (~29-38 Ma) and uplifted
(inferred) into the Miocene (apatite fission track; Coutand et al., 2001). To the northwest
and west lies the Cerro Negro, and to the south-southwest, the Sierra Alto Huasi and
Cerro Durazno ranges, all uplifted during an older phase of deformation (Oligocene-
Miocene), synchronous with uplift of the Chango Real and Alto Grande (inferred from
apatite fission track; Carrapa et al., in press; Allmendinger, 1986). To the southeast of
the basin lies the mid-late Miocene volcanic Farallón Negro (Aceñolaza et al., 1982).
The El Cajon-Campo del Arenal Basin lies approximately 100 km northeast of the
Fiambalá Basin and is bounded by the uplifted Sierra de Chango Real to the west, the
Sierra de Quilmes anticline to the east, and the uplifted Sierra de Aconquija to the
southeast. East and west-dipping high-angle reverse faults separate the basin from the
Aconquija and Chango Real ranges, respectively. Exhumation of the Chango Real commenced in the late Eocene-early Oligocene, ~29-38 Ma (apatite fission track;
Coutand et al., 2001), while active exhumation was occurring by the late Miocene, ~6
Ma, in the Aconquija range (apatite fission track; Sobel and Strecker, 2003). The Sierra de Quilmes anticline was inferred to have uplifted and exhumed prior to 7-6 Ma (apatite fission track; Mortimer et al., 2006).
26 The Santa María Basin is bordered by the Sierra de Quilmes anticline to the west and the Cumbres Calchaquíes and Sierra Aconquija ranges to the east. The basin is bounded by east-dipping high-angle reverse faults against the Cumbres Calchaquíes and Sierra
Aconquija ranges, both exhumed in the late Miocene, between ~9-5 Ma and by ~6 Ma, respectively (apatite fission track; Mortimer et al., 2006; Sobel and Strecker, 2003).
The Angastaco Basin is bounded to the west by the high-angle reverse faults of the proximal Cerro Negro range and the distal Eastern Cordillera (Cumbres Luracatao); to the east lies the Sierra de los Colorados. Exhumation of the Eastern Cordillera began in the Oligocene-Miocene (25-15 Ma), and associated deformation and uplift continued into the Pliocene (apatite fission track; Coutand et al., 2006; Deeken et al., 2006). This event was followed by inferred uplift of the Sierra de los Colorados (Santa Barbara ranges) in the early Pliocene (5.4-2.4 Ma) (apatite fission track; Coutand et al., 2006).
The Quebrada del Toro Basin is located between the Puna Plateau to the west and the
Sierra Pasha and Sierra Pasha Sur to east. The basin is separated from the plateau by a west-dipping reverse fault and from the Sierra Pasha by an east-dipping reverse fault.
Exhumation of the Eastern Cordillera occurred during the Oligocene-Miocene and inferred deformation and uplift continued into the Pliocene (apatite fission track; Coutand et al., 2006; Deeken et al., 2006). This was followed by inferred late Miocene (8-6 Ma) exhumation and uplift of the Sierra Pasha (inferred from transition of sediment depositional environments; Hilley and Strecker, 2005).
Deposition of the Punaschotter Conglomerates could have been the product of local tectonic deformation caused by uplift of basin-bounding ranges along the Puna Plateau
27 margin. If this was the case, timing of deposition of the unit should be concurrent with
local uplift, as shown in Figure 7, and will vary among basins.
2.3 CLIMATIC SETTING
2.3.1 Global Cenozoic Climate Change
High-resolution deep sea oxygen (δ18O) and carbon (δ13C) isotope records record
global climate change throughout the Cenozoic, as described by Zachos et al. (2001).
From the mid Paleocene (59 Ma) to the early Eocene (52 Ma), a pronounced warming trend occurred, culminating in the early Eocene Climatic Optimum (52-50 Ma). Cooling ensued from the early-mid Eocene (50-48 Ma) through the early Oligocene (35-34 Ma), after which time ice sheet expansion in Antarctica began and continued through the late
Oligocene (26-27 Ma). A long trend of warming followed, culminating in the mid-
Miocene climatic optimum (17-15 Ma). Subsequently, temperatures decreased as the
Antarctic ice sheet grew and sea level fell rapidly, continuing into the early Pliocene (6
Ma) (e.g. Flower and Kennett, 1995; Theide and Vorren, 1994). Warming ensued until
~3-4 Ma (Flower and Kennett, 1995) when worldwide climate conditions transitioned from warm and humid to cold and dry (Peizhen et al., 2001). The onset of climate change at ~3-4 Ma is considered the beginning of the Neogene to Quaternary glaciation, and was accompanied by a worldwide transition from a moderate and fairly unchanging climate to highly variable and stormy conditions (Donnelly, 1982; Zhang et al., 2001).
Northern Hemisphere Glaciation intensified at ~2.4 Ma (Maslin et al., 1998; Shackleton et al., 1988) and, after this time, Plio-Pleistocene glaciations were strongly dictated by
28 Global Climate Age Reference Cause Reference decreasing temperatures following~15-6 Ma Zachos et al., 2001; increase of Arctic and Antarctic Kennet and Barker, 1990 Mid-Miocene climatic optimum Flower and Kennett, 1995; ice sheets Theide and Vorren, 1994 Theide and Vorren, 1994 warm and humid ~6-3 Ma Flower and Kennett, 1995 ice sheet, orbital fluctuations Poore and Sloan, 1996 cold and dry ~3 into Pleistocene Peizhen et al., 2001 Quaternary glaciation Shackleton et al., 1988 Maslin et al., 1998
Regional increased aridity since Mid-Miocene~11-7 Ma Hartley and Chong, 2002 fluctuations in the proto-Humboldt Current Dunai et al., 2005 (South America)climatic optimum in Atacama Desert~14-9 Ma Alpers and Brimhall, 1988 uplift of the Andes Lamb and Davis, 2003 Hinojosa, L.F., 2003 cooling temperatures (Chile)Mid-Miocene to ~6.4 Ma Kött et al., 1995; increases in the Hadley circulation, cold Hartley, 2003 semi-arid Gaupp et al., 1999 Humboldt Current upwelling, and establishment arid ~6.4-3.7 Ma Kött et al., 1995; of the Panama land bridge after 4 Ma Gaupp et al., 1999 hyperaridity ~2.6 into Pleistocene Hartley, 2003
Local (Puna Margin) 29 Angastaco Basin arid ~15-9 ±1 Ma Coutand et al., 2006 regional climate during this period Strecker et al., 2007 humid ~9-5 Ma Eastern Cordillera rise, basin receives Atlantic precipitationStark and Anzóegui, 2001 Strecker, et al., 2007 semi-arid ~3.4-2.4 uplift of eastern bounding ranges Sierras de los ColoradosCoutand et al., 2006 Santa Maria Basin dry, arid ~12-9 Ma Kleinert and Strecker, 2001 regional climate during this period Strecker et al., 2007 increasing moisture ~9-7 Ma steepening relieft near eastern basin margin Kleinert and Strecker, 2001 wet-dry seasonality ~7-5 Ma uplift of the Sierra Aconquija and Cumbres Calchaquíes to theSobel east and Strecker, 2003 increased precipitation ~5-3 Ma uplift of the Sierra de Quilmes range to the west Mortimer et al., 2006 arid ~3-2.5 Ma uplift of Eastern bounding ranges Kleinert and Strecker, 2001 semi-arid ~2.5-Pleistocene windward flanks receive high precipitation Kleinert and Strecker, 2001 Calchaqui Valley arid ~15-10 Stark and Anzóegui, 2001 Pacific-driven precipitation blocked by western rangesStark and Anzóegui,01 20 humid ~10-3 Ma atmospheric circulation patterns change, now Atlantic-derivedStark precip and Anzóegui, 2001 Angastaco Basin increased aridity ~9-5 Ma Carrapa and Mortimer, in preparation eastward migration of Andes leading to basinn reorganizatioCarrapa and Mortimer, in preparation Evidence Fiambala Basin arid ~8.2-5.2 Carrapa et al., 2006, in press coarsening upward of Tamberia and Guanchin FormationsCarrapa et al.,006, 2 in press due to eastward fault propagation presence of mudcracks, halite and gypsum layers,s no plant fossilBossi et al., 1999
Table 1. Climate Summary - Past 10 Ma orbital cycles operating on of periods of 20,000-100,000 years (Hays et al., 1976). A
summary of global climate change in the past 10 Ma is shown in Table 1.
2.3.2 Orbital Forcing of Climate
Glacial and interglacial periods are intimately linked to Earth’s orbital variations, or
Milankovitch Cycles, operating on 20,000-400,000 year time scales (Hayes et al, 1976).
Eccentricity, obliquity, and precession affect the distribution, intensity, and location of solar radiation received on Earth. Variations in these parameters produce changes in the
solar radiation reaching earth, particularly on insolation amount and timing in high
northern latitudes. Thus, atmospheric moisture, albedo, snowfall, ice sheet abundance,
and seasonal contrast all have a significant affect on the warming and cooling of the
planet. At the maximum aphelion, when the Earth is farthest from the Sun, 20-30% less
radiation is received (e.g. Berger, 1988; Schwarzacher, 1993) allowing for a significant
drop in temperature, enhancing glacial advances. At a low axial tilt angle, the insolation
difference between summer and winter is small, but the insolation difference between the
Equator and the polar regions is great. Ice sheet growth is enhanced because warmer
winters promote greater amounts of moisture in the air, increasing snowfall; cooler
summers reduce snow melt, further lowering temperatures (Berger et al., 1989;
Schwarzacher, 1993). Axial wobble increases seasonality as a result of the Earth’s
proximity to the Sun and affects the distribution of solar radiation to climatic belts
(Schwarzacher, 1993). When the axis is tilted towards Vega, the Northern Hemisphere
experiences winter at the Earth’s aphelion, allowing for the maximum cold temperatures
(Schwarzacher, 1993). Orbital parameters alone will not produce a glacial period, even if
30 all the parameters favor glaciations; a positive feedback loop must also exist between
snow and ice cover, albedo, cooling atmosphere, and glacial and ice sheet growth (e.g.
Ruddiman and Prell, 1997; Berger et al., 1989).
During the late Pliocene-Pleistocene, glacial and interglacial periods have been
strongly linked to changes in orbital variations, namely increased amplification and
increased frequency of variation of orbital parameters (e.g. Zachos et al., 2001; Molnar,
2004; Hays et al., 1976). During the Triassic- Jurassic and Oligocene-Miocene, the
dominant cycle acted on a ~400 kyr cycle, implicating eccentricity as the dominant
orbital forcing (Zachos et al., 2001). At ~5 Ma (early Pliocene) the dominant cycle
switched to 41 kyr (obliquity), remaining strong until ~1.5 Ma. Since ~1.5 Ma (early
Pleistocene), 100 kyr and 20 kyr cycles, eccentricity and precession, respectively, have
strengthened (Mix et al., 1995; Pisias et al., 1995).
Increased erosion and sedimentation rates are produced in three ways: 1) increased
amplitude and increased frequency of variation of orbital parameters on 20-100 kyr
cycles, 2) increased glaciations, and 3) lowered sea level. Increased amplitude and
increased frequency of variation of orbital parameters drive climate change, forcing the
landscape to evolve and equilibrate from one state to another (Molnar, 2004). During
stable climates that varied slowly (400 kyr cycles), and in the absence of other factors
that affected erosion rates, developing landscapes and geomorphic processes were able to
equilibrate with the slow changes and little erosion resulted (Molnar, 2004). At ~4 Ma,
when climate oscillations began to fluctuate between extremes on much shorter time
periods (20,000-41,000 years), the landscape was not able to reach equilibrium and rates
of erosion increased. High-frequency forcing will result in more erosion than will low-
31 frequency forcing (Molnar, 2004). If erosion depends monotonically on climatic forcings,
then both the increased amplitude and the increased frequency of variation during the
Cenozoic, particularly in the last 2-4 Ma, should have contributed to increased erosion
and sediment accumulation (Molnar, 2004).
Glacial periods and lowered sea level also contribute to increased erosion and
sedimentation rates. Regions at high latitudes and high altitudes were covered by glaciers
since global cooling accelerated at ~3-4 Ma, which have the ability to erode and transport
large volumes of sediment. Particularly at high altitudes, glacial erosion almost surely increased when global cooling occurred during the Pliocene (Molnar, 2004). As glaciers expanded, sea levels fell, thus exposing the continental shelf, once covered by the ocean, to erosion. Hay et al. (1988) attributes the increase in sediment accumulation rates in the past 4-5 Ma to lowered sea levels and erosion of newly exposed continental margins.
Deposition of the Punaschotter Conglomerates may have been caused by increased
erosion and sedimentation rates, resulting from disequilibrium in the landscape and
increased glaciation, driven by increased orbital variations. If this is the case, deposition
of the Punaschotter conglomerate should have been regionally synchronous, and should have commenced at 2-4 Ma.
2.3.3 Continental-Scale Tectonic Forcing of Climate
Tectonic controls on climate change include tectonic uplift, openings and closings of oceanic gateways, ocean bathymetry, and the position of the continents (Crowley and
Burke, 1998). These “boundary conditions” change gradually, over millions of years, and operate unidirectionally (Zachos et al., 2001). As such tectonic changes occur, major
32 shifts can be generated in the global climate system and in the response to orbital forcings
(e.g. Zachos et al., 2001; Raymo and Ruddiman, 1992; Ruddiman and Prell, 1997).
Tectonic uplift can drive climate change in several ways: (1) uplift at temperate latitudes extends the length of winter in that region, expanding snow cover and increasing albedo, driving down temperatures (Birchfield and Weeterman, 1983; Ruddiman and
Prell, 1997); (2) General Circulation Model sensitivity tests (Raymo and Ruddiman,
1992; Ruddiman and Kutzbach, 1989) show that elevation increases in large-areas (i.e. the Tibetan and Puna-Altiplano Plateaus) can alter atmospheric circulation patterns and drive down global temperatures; and (3) physical erosion increases in actively uplifting mountains, causing chemical weathering of silicate minerals to increase, drawing down carbon dioxide through chemical reactions (e.g. Molnar and England, 1990; Raymo et al.,
1988) as rapid chemical weathering occurs. Global cooling and regional uplift have been closely linked in the Cenozoic era (e.g. Raymo and Ruddiman, 1992). As discussed below, in the Puna, plateau uplift may have caused aridification
The opening and closing of oceanic gateways can affect ocean heat transport and may
be responsible for relatively rapid climate change (e.g. Berger et al., 1981; Mikolajewicz
et al., 1993). For example, the closing of the Central American passage between the
Atlantic and Pacific at 3.5-3 Ma (Coates et al., 1993) has been credited as a possible
underlying cause of the Northern Hemisphere glaciation (e.g. Hay, 1996; Coates et al.,
1993) and increased aridity in western South America due to intensification in upwelling and expansion of the Humboldt Current, which blocks moist air masses from reaching the coast (Ibaraki, 1997). Changes in aridity through the Cenozoic in the Puna plateau have been linked to changes in the Humboldt Current, as discussed below.
33 Increased aridity is thought to influence erosion, sedimentation, and clastic grain size
in a number of ways. In some climates, increased precipitation can lead to increased erosion; in other climates, a decrease in precipitation, due to increased aridity, can lead to
a lack of vegetation as the result of decreased water availability (Hall, 1990). Vegetation
protects the surface from erosion, and a decrease in vegetation can result in increased
erosion and sediment transport. Baroclinic instability resulting from climate change can
lead to a stormier climate (Molnar and England, 1990), whereby the magnitude and/or
frequency of large floods are increased. Arid regions tend to be affected by a larger range
of flood magnitudes than do more humid regions, and rare floods are often of a larger magnitude than are annual floods (Pitlick, 1994). Additionally, the frequency of floods can also be increased in arid regions, where more large-magnitude floods occur per small-magnitude flood (Molnar, 2001). Despite the fact that increased aridity results in less precipitation and river discharge, larger floods have a greater duration and can carry a larger amount and size of sediment than can smaller floods. Thus, large, yet infrequent floods, may be more effective erosional agents and have the heightened energy to carry large cobbles (Molnar, 2001).
As the climate oscillates between glacial and interglacial states, frost action fractures
the surface rock, transforms the rock to debris, and transports it to valleys through glacial
erosion. Increased climate variability might accelerate erosion and increase clastic grain
size through mass wasting processes (Peizhen et al., 2001).
Aridity in the South American Andes has been a long-lived climatic feature, and may
have been established as early as the Mesozoic (Hartley et al., 1992). The current
atmospheric and oceanic circulation system that drives cold upwelling of the Humboldt
34 Current, low rates of evaporation, and low precipitation along the west coast of the continent may have been established since the Paleogene (e.g. Parrish et al., 1982).
Sediments in the western and central Andes record semi-arid to arid conditions during the
Cretaceous and Eocene-Oligocene (Adelmann, 2001; Voss, 2002). More recent aridification is reported to have occurred during the mid-Miocene climatic optimum at
~11-7 Ma (Hartley and Chong, 2002) and ~14-9 Ma (Alpers and Brimhall, 1988) in the
Atacama Desert, during which time conditions changed from semi-arid to hyper-arid and temperatures fell (Hinojosa, L.F., 2003). This is largely thought to have been the result of fluctuations in the proto-Humboldt Current (Dunai et al., 2005) and uplift of the Andes
(Lamb and Davis, 2003). Other data suggest a semi-arid climate until 6.4 Ma in Chile, transitioning to aridity from 6.4-3.7 Ma (Kött et al., 1995; Gaupp et al., 1999) due to continued uplift of the Andes and changes in the Humboldt Current. After 2.6 Ma, hyper- arid conditions were dominant along the west coast of South America. This transition would have likely been the result of further intensification of the Hadley circulation, the cold Humboldt Current, and the establishment of the Panama land bridge after 3-3.5 Ma
(e.g. Hartley, 2003).
Deposition of the Punaschotter Conglomerates could have been the result of increased aridity due to regional climate change, enhanced by changing boundary conditions like plateau uplift and the closing of the Central American Seaway. If this is the case, deposition should align with one of two events: 1) increased deposition should be documented between 14-7 Ma (e.g. Hartley and Chong, 2002; Alpers and Brimhall,
1988) as the result of a climate transition from semi-aridity to hyper-aridity, due to proto-
Humboldt Current fluctuations and uplift of the Andes, or 2) increased aridity from 6.4
35 through 2.6 Ma (e.g. Kött et al., 1995; Gaupp et al., 1999; Hartley, 2003) driven by continuing Andean uplift, changes in the Humboldt Current and Hadley circulation, and, after 3-3.5 Ma, closing of the Central American Passage. The latter possibility aligns more closely with global climate change recorded between 2-4 Ma, thus is a more probable cause for deposition of the Punaschotter Conglomerates. Regardless, either climate event should have caused regionally synchronous deposition along the eastern
Puna margin.
2.3.4 Local Tectonic Forcing of Climate
Mountain range uplift has clear and direct impacts on atmospheric circulation and precipitation patterns in the associated region. As a mountain range is uplifted, an orographic barrier is created, resulting in increased precipitation on the windward side and decreased precipitation or increased aridity on the leeward side.
In the Puna Plateau, aridification has typically been linked to the formation of internal drainage and the deposition of halite and gypsum-bearing units between 24-15 Ma
(Alonso et al., 1991; Vandervoort et al., 1995).
In the Angastaco basin, Coutand et al (2006) show that, from 15-9 Ma, the Angastaco
Formation was deposited under arid conditions, in contrast to the Palo Pintado Formation, deposited between 9-5 Ma under a more humid climate. By 9-5 Ma, the Eastern
Cordillera to the west had reached high elevation, blocking Atlantic-derived moisture and depositing it on the windward Angastaco Basin. In the late Pliocene (3.4-2.4 Ma) an orographic barrier was built east of the basin, cutting it off from precipitation and establishing semi-arid conditions (Coutand et al., 2006).
36 Stark and Anzóegui (2001) document a climate change in the late Miocene, from aridity (~15-10 Ma, upper Angastaco Formation) to humidity (~10-3 Ma, lower Palo
Pintado Formation), but they attribute these changes to variations in atmospheric circulation patterns over South America, not to uplift of orographic barriers. When the
Angastaco Formation was deposited (pre-10 Ma), a mountain range existed to the west, placing the Angastaco basin in the rain shadow, starving it of precipitation from the west
(Pacific). Between 10-9 Ma, winds from the Pacific switched to Atlantic-derived winds, due to variations in atmospheric circulation patterns, which effectively altered the rain shadow location (Stark and Anzóegui, 2001). The Angastaco Basin was then located on the windward side of the mountains, receiving greater amounts of precipitation and accounting for a more humid climate. Because topography was much the same from the late Miocene to the early Pliocene, Stark and Anzóegui (2001) argue that uplift of orographic barriers was not the likely cause of the climate change.
A recent study by Carrapa and Mortimer (in preparation) shows that the late Miocene was a time of basin and drainage reorganization in the Angastaco Basin, due to eastward migration of the tectonic orogen. They propose that the change in depositional environments from fluvial to lacustrine, attributed by Stark and Anzóegui (2001) to altered atmospheric circulation patterns, was actually the result of a decrease in sedimentation rates and constant or increasing subsidence rates driven by tectonic deformation and basin reorganization. These authors argue that increased aridity is, in fact, recorded in the sediments of intramontane basins, and not increased humidity, as suggested by Stark and Anzóegui (2001) and Stark and Vergani (2001).
37 In the Santa María Basin, paleosols with illuvial clay and organic matter, calcic and silicic rhizoliths, authigenic clays, and C3 and C4 plants are proxies for climate change since ~12 Ma (Kleinert and Strecker, 2001). A seasonally dry climate existed after ~12
Ma with increased moisture documented from ~9-7 Ma. From 7-5 Ma, wet-dry
seasonality existed, the result of uplift of the Sierra Aconquija and Cumbres Calchaquíes
ranges to the east (Sobel and Strecker, 2003). At 5-3 Ma, uplift of the Sierra de Quilmes
range to the west brought increased precipitation into the basin (e.g. Mortimer et al.,
2006). Aridification ensued from 3-2.5 Ma, the result of further uplift of eastern bounding
ranges (Kleinert and Strecker, 2001), and semi-arid conditions were established by the
end of the Pliocene.
The Tamberia and Guanchin Formations of the Fiambalá Basin, ~8.2 Ma and ~5.5
Ma, respectively, contain numerous mudcracks or mud layers, halite and gypsum layers,
and no plant fossils, suggesting an arid, or episodically dry environment by the time the
late Miocene units were deposited (Bossi et al., 1999; Carrapa et al., in press). Table 1
displays a summary of both regional and local climatic changes.
Tectonically driven uplift along the southern margin of the Puna Plateau clearly has
altered local climate through the formation of orographic barriers. If Punaschotter
deposition was a result of such changes (either an increase or decrease in humidity), the timing of deposition in each basin will be linked to local uplift and deformation, and the timing will vary from basin to basin.
38 2.4 METHODS
To more accurately constrain the depositional age of the Punaschotter Conglomerates,
interbedded ashes and ignimbrites contained within or near the conglomerates were
sampled for zircon U-Pb geochronology. U-Pb dating has proven to be a powerful tool for constraining the stratigraphic age of continental clastic sequences (e.g. Coutand et al.,
2006; Carrapa et al., in press). Eight samples were collected at the surface along the western margin of the Fiambalá Basin (Figure 8). Five ash samples and one ignimbrite sample were collected from within the Punaschotter Conglomerates, one ash sample from within the Guanchin Formation, and one sample from a thick ignimbrite unit that unconformably underlies the Punaschotter in the northern section of the basin.
Of these eight samples, four were collected from the northern region of the basin,
three from the central region, and one from the southern region, complementing previous
U-Pb dating carried out by Carrapa et al. (in press).
Detailed structural, geologic, and geomorphic mapping was conducted in the areas of
sample collection to better interpret and constrain the timing of deposition and
deformation of the Punaschotter Conglomerates along the western margin of the
Fiambalá Basin. Corresponding cross sections were constructed to further define
structures, the distribution of ash deposits, thicknesses of Miocene-Pliocene units, and
subsurface interpretation. From these cross sections, stratigraphic distances of each ash
unit above and below the Punaschotter-Guanchin contact were determined and are
illustrated in Figure 16. Mapping was carried out on orthorectified 1:20,000 air photos
and 1:20,000 Google Earth maps.
39
Figure 8. Shaded relief map of the Fiambalá Basin. Cross sections A-A’ through F-F’ with locations and ages of 6 U-Pb dated ashes and 2 U-Pb dated ignimbrites shown. Location of Rio Guanchin cross section by Carrapa et al. (in press) shown.
40 Approximately 2 kg samples were collected and prepared for analysis. Samples were
crushed and sieved, with the 125 to 75 μm fraction retained. Samples were then washed,
dried, sonicated, and washed and dried again to remove adhering small particles. The
heavy zircon fraction was separated using bromoform and methylene iodide, and a Frantz
magnet was used for the final separation.
Zircon analysis was performed at the University of California at Los Angeles
(UCLA) under the direction of Dr. Axel Schmitt. Grains were hand selected, epoxy
mounted, and polished to expose the interior of each grain. Ultrasonic cleaning was performed with soapy water, diluted HCl and distilled water. Finally, the mount was coated with ~10 nm of Au in a high vacuum chamber (>10-8 Torr) and placed in a
Cameca IMS 1270 ion microprobe for analysis. A mass-filtered, ca. 15 nA 16O- beam was
focused on a ca. 30-35 μm diameter spot to date each grain. Roughly 10 grains per
94 + 204 + sample were analyzed to ensure robust age constraints. Intensities of Zr2O , Pb ,
206Pb+, 207Pb+, 208Pb+, 238U+, 232Th16O+, and 238U16O+ were measured in 10 cycles per grain
to further ensure accurate age constraints and to eliminate partially reworked zircons.
Calibration standard zircon AS-3 (Paces and Miller, 1993) grains were used to
determine relative sensitivities of Pb and U, using calibration techniques similar to
Compston et al. (1994). Th and U contents were calculated by multiplying
232 16 + 94 + 238 16 + 94 + Th O / Zr2O and U O / Zr2O ratios of the unknowns with corresponding
relative sensitivity values determined on reference zircon 91500 (Wiedenbeck et al.,
1995). Further details of this process can be found in Schmitt et al. (2003).
41 Correlation Sample Grain 238U/206Pb 238U/206Pb 207Pb*/ 207Pb*/ of T-W Concordia 206/238 age ±1s U Th UO/U % 206Pb* 206Pb* 206Pb* Ellipses [Ma] [Ma] ppm ppm
1 s.e. 1 s.e. Mean = 3.66 ± 0.10 FA3 1 1693.2 57.6 0.0440 0.0041 -0.09 3.90 0.13 813 429 9.4 100.3 [2.6%] 95% confidence (Ignimbrite) 2 1652.9 84.1 0.1403 0.0182 0.21 3.47 0.22 470 774 9.5 88.0 Weighted by data point errors only 3 1810.3 72.8 0.0600 0.0075 0.35 3.53 0.15 817 1392 9.5 98.2 0 of 14 rejected 4 1768.7 68.2 0.0723 0.0066 0.02 3.57 0.14 742 1028 9.4 96.6 MSWD = 0.62, probability = 0.84 5 1510.6 100.6 0.1430 0.0229 0.13 3.79 0.31 154 184 9.5 87.6 6 32.1 1.0 0.8203 0.0051 -0.01 2.11 8.86 538 1181 9.2 1.0 7 234.6 8.4 0.7278 0.0154 -0.03 3.58 1.43 716 899 10.0 12.8 8 1267.1 49.9 0.2898 0.0087 -0.05 3.54 0.22 1298 2068 9.2 68.8 9 858.4 39.3 0.4555 0.0209 0.21 3.62 0.41 844 1248 9.1 47.7 10 1463.5 147.6 0.1856 0.0234 0.41 3.67 0.46 152 190 9.5 82.2 11 1630.8 61.4 0.0938 0.0113 -0.30 3.76 0.16 1109 1440 9.2 93.9 12 1696.1 58.4 0.0734 0.0066 0.15 3.72 0.13 607 830 9.2 96.5 13 1698.4 60.3 0.0647 0.0039 -0.32 3.73 0.14 1093 2048 9.3 97.6 14 1825.5 71.3 0.0644 0.0117 0.17 3.49 0.15 624 892 9.5 97.7 w.m. 3.66 ±1σ 0.05 MSWD 0.62 n 14 FA4 1 1624.2 40.6 0.0522 0.0036 0.36 4.02 0.10 1150 559 8.2 99.2 Mean = 3.99 ± 0.24 (Ash) 2 1712.0 51.6 0.0579 0.0025 -0.28 3.79 0.11 1134 674 9.1 98.5 [6.0%] 95% confidence 3 1575.8 45.7 0.0748 0.0044 0.12 4.01 0.12 1297 1041 8.5 96.3 Weighted by data point errors only 4 1523.0 50.3 0.0624 0.0056 0.20 4.18 0.14 601 944 8.1 97.9 0 of 4 rejected w.m. 3.99 MSWD = 1.7, probability = 0.17 ±1σ 0.06 MSWD 1.67 n 4 FA5 1 1465.2 48.1 0.0522 0.0033 -0.02 4.47 0.15 1104 172 9.2 99.2 Mean = 4.12 ± 0.13 42 (Ash) 2 1533.5 81.4 0.0684 0.0060 -0.12 4.18 0.23 381 74 8.7 97.2 [3.2 %] 95% confidence 3 1573.6 46.8 0.0504 0.0033 0.20 4.16 0.12 1601 622 8.9 99.5 Weighted by data point errors only 4 1470.4 51.5 0.0782 0.0065 -0.24 4.30 0.16 1104 235 9.2 95.9 0 of 11 rejected 5 1640.7 55.2 0.0675 0.0056 0.04 3.91 0.14 1338 537 9.1 97.3 MSWD = 1.8, probability = 0.060 6 1599.7 50.9 0.0561 0.0033 0.25 4.07 0.13 1615 666 9.2 98.7 7 1631.3 53.0 0.0646 0.0048 0.09 3.95 0.13 1330 340 9.1 97.6 8 1447.2 58.0 0.0563 0.0053 0.03 4.50 0.18 427 39 8.9 98.7 9 1627.3 68.1 0.0605 0.0065 -0.11 3.97 0.17 513 341 9.3 98.2 10 1626.8 55.6 0.0596 0.0051 0.03 3.97 0.14 649 423 9.3 98.3 11 1583.3 59.7 0.0574 0.0060 -0.09 4.11 0.16 758 242 8.8 98.6 w.m. 4.12 ±1σ 0.04 MSWD 1.77 n 11 FA6 1 1548.2 63.5 0.0540 0.0048 -0.22 4.22 0.17 350 110 7.7 99.0 Mean = 4.08 ± 0.09 (Ash) 2 1624.4 38.0 0.0465 0.0034 0.08 4.06 0.09 1178 499 7.6 100.0 [2.2 %] 95% confidence 3 1614.2 36.7 0.0528 0.0032 0.21 4.05 0.09 1412 573 8.1 99.1 Weighted by data point errors only 4 605.7 13.5 0.5381 0.0078 0.07 4.03 0.29 1495 777 7.6 37.1 0 of 6 rejected 5 1530.7 51.3 0.0723 0.0078 0.27 4.16 0.15 542 215 7.7 96.6 MSWD = 0.23, probability = 0.95 6 1605.4 30.7 0.0516 0.0025 -0.16 4.07 0.08 1814 1012 8.0 99.3 w.m. 4.08 ±1σ 0.05 MSWD 0.23 n 6 The ion microprobe facility at UCLA in partly supported by a grant from the Instrumentation and Facilities Program, Division of Earth Sciences, National Science Foundation.
Table 2. U-Pb Geochronology Results Correlation Sample Grain 238U/206Pb 238U/206Pb 207Pb*/ 207Pb*/ of T-W Concordia 206/238 age ±1s U Th UO/U % 206Pb* 206Pb* 206Pb* Ellipses [Ma] [Ma] ppm ppm
FA7 1 1671.1 70.7 0.0595 0.0078 -0.23 3.89 0.17 403 89 8.8 98.3 Mean = 3.93 ± 0.10 (Ash) 2 1682.1 54.3 0.0381 0.0039 0.36 3.96 0.12 1038 361 9.4 101.0 [2.3 %] 95% confidence 3 1691.2 84.1 0.0437 0.0074 0.34 3.92 0.19 433 143 8.9 100.3 Weighted by data point errors only 4 916.6 60.2 0.3787 0.0167 -0.27 4.13 0.57 941 482 8.7 57.5 0 of 13 rejected 5 1615.0 57.1 0.0938 0.0096 0.36 3.84 0.15 638 198 9.0 93.9 MSWD = 1.02, probability = 0.43 6 1628.4 57.3 0.0553 0.0048 0.12 4.01 0.14 1492 487 9.0 98.8 7 1141.4 42.9 0.2968 0.0087 -0.03 3.92 0.23 1226 577 8.8 67.9 8 336.8 10.7 0.6739 0.0083 0.01 3.86 0.80 1689 751 8.7 19.7 9 1572.3 68.2 0.0681 0.0069 -0.19 4.08 0.18 436 151 8.8 97.2 10 1624.7 62.0 0.0625 0.0048 0.08 3.97 0.15 787 442 9.0 97.9 11 1709.7 56.1 0.0627 0.0045 0.19 3.78 0.13 1095 399 9.4 97.9 12 1514.9 56.9 0.0494 0.0048 -0.04 4.32 0.16 711 440 9.2 99.6 13 1399.0 50.1 0.2345 0.0124 0.07 3.59 0.18 618 247 9.2 75.9 w.m. 3.93 ±1 0.05 MSWD 1.02 n 13 FA8 1 1205.4 81.8 0.2608 0.0570 -0.62 3.97 0.59 218 89 9.0 72.5 Mean = 4.11 ± 0.10 (Ash) 2 1510.3 52.7 0.0574 0.0041 0.05 4.28 0.15 1188 959 9.1 98.6 [2.3 %] 95% confidence 3 1615.2 62.9 0.0634 0.0043 0.17 3.98 0.16 1152 693 9.1 97.8 Weighted by data point errors only 4 1520.2 51.5 0.0519 0.0026 0.03 4.30 0.14 2372 878 9.1 99.3 0 of 12 rejected 5 1603.1 56.8 0.0574 0.0048 0.22 4.05 0.14 1220 672 9.2 98.6 MSWD = 0.97, probability = 0.47 6 1556.2 52.6 0.0569 0.0046 0.04 4.18 0.14 837 321 9.1 98.6 7 853.2 40.7 0.4238 0.0121 -0.12 3.98 0.43 786 556 9.3 51.7 8 1533.7 44.2 0.0548 0.0025 0.22 4.23 0.12 4117 3107 8.9 98.9 9 1586.0 59.6 0.0606 0.0060 0.23 4.08 0.16 536 162 9.2 98.1 10 1255.7 36.3 0.2483 0.0118 -0.11 3.90 0.17 1449 601 9.1 74.2 11 1558.4 61.9 0.1361 0.0121 -0.19 3.75 0.18 743 251 9.4 88.5
43 12 1572.6 63.1 0.0524 0.0031 0.20 4.15 0.17 1461 671 9.1 99.2 w.m. 4.11 ±1 0.05 MSWD 0.97 n 12 FA9 1 1618.6 70.0 0.0639 0.0063 0.11 3.93 0.17 602 956 8.7 97.7 Mean = 3.90 ± 0.12 (Ash) 2 1587.3 74.8 0.0676 0.0074 0.08 4.00 0.20 616 813 8.9 97.3 [3.0 %] 95% confidence 3 1554.7 67.9 0.0523 0.0039 0.00 4.15 0.18 876 1433 8.5 99.2 Weighted by data point errors only 4 1627.3 59.3 0.0858 0.0049 -0.15 3.81 0.15 1180 1415 8.8 94.9 0 of 9 rejected 5 1653.4 81.2 0.0939 0.0107 0.08 3.70 0.20 550 828 8.6 93.9 MSWD = 1.06, probability = 0.39 6 1564.0 74.4 0.1259 0.0110 0.39 3.76 0.20 379 460 9.0 89.8 7 1731.3 120.5 0.0780 0.0104 0.30 3.64 0.26 235 216 9.0 95.9 8 1699.5 73.1 0.0636 0.0058 0.07 3.76 0.17 781 992 8.9 97.8 9 1539.9 55.0 0.0689 0.0039 0.16 4.15 0.15 2122 960 8.9 97.1 w.m. 3.90 ±1 0.06 MSWD 1.06 n 9 Ignimbrite 1 892.9 38.0 0.0620 0.0058 0.33 7.15 0.31 398 257 9.2 98.0 Mean = 7.59 ± 0.25 2 797.4 29.6 0.0515 0.0034 0.10 8.12 0.30 844 318 8.8 99.3 [3.3 %] 95% confidence 3 825.1 41.8 0.0542 0.0052 0.37 7.81 0.40 339 225 8.9 99.0 Weighted by data point errors only 4 939.0 53.5 0.0671 0.0072 0.19 6.76 0.40 240 131 8.6 97.3 0 of 10 rejected 5 829.2 32.6 0.0563 0.0043 0.06 7.74 0.31 584 490 9.2 98.7 MSWD = 1.2, probability = 0.27 6 829.2 29.9 0.0491 0.0034 -0.02 7.83 0.28 1801 733 8.9 99.6 7 811.7 40.1 0.0691 0.0046 -0.04 7.77 0.40 425 411 8.8 97.1 8 856.9 32.6 0.0503 0.0027 0.19 7.55 0.29 1471 1221 9.1 99.5 9 854.7 28.3 0.0587 0.0027 0.03 7.50 0.25 1579 846 9.3 98.4 10 850.3 32.3 0.0608 0.0036 -0.04 7.52 0.29 767 368 9.2 98.1 w.m. 7.59 ±1 0.10 MSWD 1.23 n 9 The ion microprobe facility at UCLA in partly supported by a grant from the Instrumentation and Facilities Program, Division of Earth Sciences, National Science Foundation.
Table 2. U-Pb Geochronology Results 2.5 RESULTS
The locations and ages of six ash samples and two ignimbrite samples, in addition to those samples collected by Carrapa et al. (in press), are shown in Figures 9-15. Figures
9a-14a depict the six geologic maps constructed for the study area, accompanied by six geologic cross sections, Figures 9b-14b. Additional mapping by Carrapa et al. (in press) includes the location of ash sample FA-9 (Figures 15a-15b). Locations and ages are also displayed in Figure 8, and the geochronologic U-Pb results are presented in Table 2.
Sample FA-3 (27° 06’ 57.6” S and 67° 49’ 03.7” W) was collected from an ignimbrite ~5 m thick, in the upper portion of the Punaschotter Conglomerates approximately 600 m above the base of the unit, east of the Precambrian-Cambrian basement contact (Figures 9a & b). Fourteen total grains were analyzed, giving an age of
3.66 ± 0.10 Ma. This is the youngest age dated for the Punaschotter Formation and, combined with three ash samples collected by Carrapa et al. (in press), likely represents a more recent eruption than do the remainder of the ash samples. Cross section A-A’
(Figure 9b) shows that this ignimbrite unit lies near the top of the Punaschotter Formation in the northern section of the Fiambalá Basin. It was likely sourced from the west, as evidenced by the inclusion of Permian lithologic clasts in the ignimbrite, because the
Permian lithology is exposed only to the west. Here the underlying Guanchin and
Tamberia Formations are not exposed. Cross section A-A’ interprets the Guanchin
Formation to unconformably lie above the lower ignimbrite unit and conformably below the Punaschotter Formation, but to pinch out in the subsurface against the ignimbrite.
This interpretation is based on the fact that the Guanchin Formation crops out to the south
(Figure 10a), where it conformably underlies the Punaschotter Conglomerates.
44
Figure 9. Map and cross section A-A’ for ignimbrite samples FA-3 and Ign shown. Portions of mapping and interpretation borrowed from Schoenbohm (with permission).
45 Elsewhere, a thick unit of ignimbrite unconformably underlies the Punaschotter
Formation and a ten-grain zircon analysis yields an age of 7.59 ± 0.25 Ma for that
ignimbrite. The age and deposition of this ignimbrite unit are not discussed further in this
analysis of Punaschotter deposition since it is an older unit underlying the conglomerates
and does not generally aid in the age constraints for deposition.
Sample FA-4 (27° 08’ 45.4” S and 67° 48’ 01.9” W) was collected from the mid-to-
upper portion of the Punaschotter Formation, 500 m above the base of the unit (Figures
10a & b). Only four grains were analyzed, yielding an age of 3.99 ± 0.24 Ma. Sample
FA-5 (27° 08’ 55.8” S and 67° 47’ 47.0” W) is an ash collected from the mid-lower
portion of the Punaschotter Conglomerates near the site of sample FA-4 and is
approximately 450 m above the Punaschotter-Guanchin transition (Figure 10b). Eleven
grains from sample FA-5 were analyzed, giving an age of 4.12 ± 0.13 Ma. Given the
overlap in ages of these samples, proximity of their locations, and their similar
stratigraphic positions, FA-4 and FA-5 may be the same ash collected from different locations. For further discussion we therefore use a weighted average of 4.09 ± 0.11 Ma for both samples, as shown in cross section B-B’ (Figure 10b). Both samples have been significantly reworked and represent the oldest age of deposition of the Punaschotter
Conglomerates in the Fiambalá Basin.
Figure 11 is a transect to the south of sample locations FA-4 and FA-5 but does not
contain a sample site. Here, the Guanchin Formation is transitional with the overlying
Punaschotter Formation. This transect will be discussed under the “deposition, folding,
and faulting” section.
46
Figure 10. Map B-B’ shows sample locations of both FA-4 and FA-5, while cross section B-B’ shows the combined location and age of both samples. Portions of mapping borrowed from Schoenbohm (with permission).
47
Figure 11. Map and cross section for transect in the northern Fiambala Basin. No ash sample was collected along this transect. Thickness of Guanchin Formation based on measurements by Carrapa et al. (in press).
48 Sample FA-6 (27° 21’ 53.7” S and 67° 52’ 24.1” W) is the only sample collected
within the Guanchin Formation (Figures 12a & 12b). It is not possible to determine the
depth of sample FA-6 below the Punaschotter-Guanchin contact because the strata
containing the ash are within a structurally bounded block. Six grains were analyzed,
yielding the age of 4.08 ± 0.09 Ma. We expected that this sample would be the oldest
dated in our study, since all other samples were collected from the overlying Punaschotter
Formation. However, the age of sample FA-6 is younger than two other samples collected from within the Punaschotter Conglomerates.
Sample FA-7 (27° 18’ 39.6” S and 67° 52’ 17.9” W) was collected within the
Punaschotter Formation, approximately 50 m above its transitional contact with the
Guanchin Formation (Figures 13a & b). Thirteen analyzed grains yielded an age of 3.93 ±
0.10 Ma. Figure 13a also identifies the locations of two samples collected by Carrapa et
al. (in press) along the Northern transect.
Sample FA-8 (27° 32’ 28.0” S and 67° 50’ 02.1” W) was collected from the upper
portion of the Punaschotter Formation exposed in the central part of the basin,
approximately 100 m above the base of the formation (Figures 14a & b). In the center of
the transect, much of the Punaschotter has been eroded away and lies unconformably atop
the Tamberia Formation. At the western and eastern margins of the transect, the
Punaschotter Formation unconformably overlies the Guanchin Formation. An age of 4.11
± 0.10 Ma was obtained from twelve analyzed zircon grains. The collected ash sample
was taken from the uppermost of four closely spaced ashes found along this transect.
Sample FA-9 (27° 42’ 31.2” S and 67° 48’ 35.1” W) was collected from the
Punaschotter Formation south of the Rio Guanchin in the southern section of the
49
Figure 12. Map and cross section E-E’ for ash sample FA-6 shown. Portions of mapping borrowed from Hauer (Carrapa et al., in press). Thickness of Guanchin and Tamberia Formations based on measurements by Carrapa et al. (in press).
50
Figure 13. Map and cross section D-D’ for ash sample FA-7. Portions of mapping borrowed from Hauer and Carrapa (Carrapa et al., in press). Locations of samples collected from within Guanchin (J28/04) and Punaschotter (J04/04) Formations are indicated. Thickness of Guanchin and Tamberia Formations based on measurements by Carrapa et al. (in press). 51
Figure 14. Map and cross section F-F’ for sample FA-8 shown. Thickness of Tamberia Formation based on measurements by Carrapa et al. (in press).
52
Figure 15. Rio Guanchin transect showing mapping and sample locations by Carrapa et al. (in press). Location of sample FA-9 from this study also shown on map and cross section.
53 Fiambalá Basin, approximately 100 m above the contact with the Guanchin Formation
(Figures 15a & b). Nine zircon grains yielded an age of 3.90 ± 0.12 Ma. This age is
slightly older than the age of the three ashes collected by Carrapa et al. (in press) on the
north side of the river.
2.6 DISCUSSION
2.6.1 Age of the Punaschotter Conglomerate in the Fiambalá Basin
Collectively, ignimbrite sample FA-3 (3.66 ± 0.10 Ma) and the three ashes collected along the Guanchin transect by Carrapa et al. (in press) (3.69 ± 0.06 Ma, 3.74 ± 0.05 Ma, and 3.77 ± 0.05 Ma, from top to bottom) likely are the same age and therefore may be from the same eruption. The weighted average of these four samples is 3.73 ± 0.03 Ma.
This age represents a more recent eruption than that of the other samples. The four above samples were collected near the top of the Punaschotter Formation, though it is unknown how much of the unit has been eroded. Based on an age of 3.05 ± 0.44 Ma for material collected from the Northern transect by Carrapa et al. (in press), deposition of the
Punaschotter Conglomerates must have continued through at least 3.05 Ma.
The remaining six samples are within uncertainty of each other and could be from the
same eruption. The weighted average of samples FA-4, -5, -6, -7, -8, and -9 is 4.02 ± 0.04
Ma. Five of the samples were collected above the Punaschotter-Guanchin contact, and one below. Two interpretations exist to explain this relationship.
The first interpretation assumes that these six samples represent two or more ashes
within the Fiambalá Basin; the ash in the stratigraphically lower Guanchin Formation
could be older, but it is difficult to discriminate this, given the uncertainty in our dating.
54 Geochemical data from these ashes would be useful for determining if these six ashes
record a single eruption, or represent multiple events, and also for explaining the younger
age of FA-6. Despite the younger age of sample FA-6, the transition between the
Guanchin and Punaschotter Formations is close to the average age of the 6 ashes, 4.02 ±
0.04 Ma.
The second interpretation is that these six ash samples represent the same eruption but
that the age of the Guanchin-Punaschotter transition varies along strike within the basin;
the transition is younger where the ash was collected in the Guanchin Formation. Further
work beyond the scope of this study is necessary to definitively assess this variation.
However, we explore the implications of this interpretation briefly here. Figure 16 plots
the distribution of all dated ash ages (both above and below the Punaschotter-Guanchin contact) versus along strike location in the basin. The approximate locations and appropriate along-strike distances between samples are shown from south to north and
the stratigraphic distances above or below the contact were measured from the cross
sections (Figures 9a-15a). Age in Ma is indicated to the left and elevations of samples
above or below the contact, and of the contact itself, are indicated to the right. Ages of
samples collected along the Rio Guanchin and Northern transects by Carrapa et al. (in
press) are also displayed for comparison. The location of the Punaschotter-Guanchin
contact is approximate, measured from generated cross sections, and does not strictly
represent the age of the contact along strike, as sedimentation rates may have varied.
However, the key point depicted is that there appears to be no along-strike correlation
between the ages from south to north, nor does there appear to be any along-strike pattern
for the contact of the Punaschotter-Guanchin. If the six samples from this study do
55
Figure 16. Distribution of dated samples of the Punaschotter Conglomerates vs. along strike location. Approximate locations of each sample from south to north shown in black. Ages of ashes sampled along the Rio Guanchin and Northern transects by Carrapa et al. (in press) shown in blue. Contact of Guanchin and Punaschotter Formations shown by red line and represents elevation, in kilometers, of the contact. Approximate stratigraphic distances of each sample above or below the contact were measured from constructed cross sections. Age in Ma shown on the right, elevation in km shown on the left.
56 represent the same eruption, then the Punaschotter-Guanchin transition varies in age along strike. It was not possible to collect ash along each river transect throughout the western margin of the basin, and large gaps exist between FA-4&5 and FA-7, between
FA-6 and FA-8, and between FA-8 and FA-9. If samples were collected and dated from these areas, a stronger interpretation could be made regarding variation of the age of the contact along strike.
The Fiambalá Basin is small enough that it should not have been susceptible to
differing effects of local climate throughout the basin and so the variation in age of the
Punaschotter-Guanchin contact must be the result of localized tectonics, short-term
events, or differences in paleowatershed size.
Uplift of bounding ranges may have been the cause of Punaschotter deposition, by
providing a sediment source, and small differences in timing of uplift could be reflected
in the along-strike age variations. Figure 16 indicates that the oldest age of the
Punaschotter-Guanchin contact at samples FA-4, FA-5, and FA-8 may be as old as 4.2
Ma. The age of the contact at samples FA-7 and FA-9 is ~3.95 Ma. If sample FA-3 is
taken into account (though deposition is likely from a different eruption), the contact age
could be as young as ~3.82 Ma in the northern end of the basin. Differences in timing of
bounding range uplift over an interval of ~0.4 My could have produced local differences
in deposition of this unit. However, our mapping suggests that movement along a single
main fault has driven local deformation, and variation in uplift rates along this single
structure may not have been large enough to account for localized uplifts that produce
this variation in the age of the contact.
57 Localized uplift could also have been coupled with, or independent of, large short-
term depositional events, such as landslides, floods, and earthquakes. Along-strike
variation in the onset of deposition could have resulted from localized inputs, such as
large floods or landslides that supplied sediment into one area, while other areas
remained unaffected. An event or series of events must have driven this variation, such as
abnormally high amounts of precipitation or glacial melt causing flooding, or an
earthquake or failure of a hillslope resulting in a landslide. If abundant precipitation resulted in infrequent, yet large-magnitude floods due to increased aridity (Molnar,
2001), higher rates of sediment accumulation should correlate with larger watersheds. If glacial melt was the case, then climate warming or a volcanic eruption should align with the age of the Punaschotter-Guanchin contact. Warming is not likely the culprit, since an increasingly arid climate is generally associated with decreasing temperatures. A volcanic eruption should be documented through ash layers in the sediment above the contact.
However, sample FA-06 was collected from below the Punaschotter-Guanchin contact. A large-magnitude flooding event, or period of large-magnitude floods, could have enhanced precipitation and increased localized sedimentation, resulting in variation in the
Punaschotter-Guanchin contact.
Additionally, rupture along individual faults could have been a cause of locally rapid
sediment deposition, resulting in areas of older deposits of conglomerates. Evidence of
localized fault rupture and fault offset would be needed to more accurately constrain
variation along-strike and these events would have had to occur before ash layers were
deposited.
58 We also hypothesize that differences in the watershed size of various river transects could be responsible for variation in the Punaschotter-Guanchin contact along strike: larger watershed would correspond to older depositional ages. Here, we assume that paleodrainage is the same as present drainage, since we are unable to completely reconstruct the paleodrainage in the basin. Figure 17 clearly shows that the paleowatershed associated with sample FA-9 is much larger than any of the remaining paleowatersheds. The next largest paleowatershed is that associated with sample FA-6, followed by FA-4 & 5, and FA-3. The two smallest paleowatersheds are those associated with samples FA-8 and FA- 7. The ages of these samples yield no pattern in comparison to paleowatershed size. FA-8 is the oldest sample yet belongs to one of the two smallest paleowatersheds. FA-6, the third oldest sample, comes from the second largest paleowatershed, while FA-4 & 5, the second oldest sample, comes from the third largest paleowatershed. FA-7, the fourth oldest sample, comes from one of the smallest paleowatersheds, while FA-9, the second youngest sample, is from the largest paleowatershed. FA-3, though grouped separately from the other samples, is the youngest sample yet is associated with a moderate-sized paleowatershed. The paleowatersheds in
Figure 17 are a drawn estimation of actual size, based on topography and present
drainage patterns, but illustrate no clear correlation between size and age, thus, are not
likely to have caused variation of the Punaschotter-Guanchin contact along strike.
However, a more accurate analysis of paleowatershed size (i.e. computer modeling)
should be made to definitively assess this interpretation.
Regardless of whether the six samples, FA 4-9, represent single or multiple events,
they tightly bracket the onset of Punaschotter deposition to 4.02 ± 0.04 Ma. This is a
59
Figure 17. Approximate size of watersheds along the western margin of the Fiambala Basin outlined in red. Sample locations are indicated in white. Dotted purple line represents known paleodrainage boundary and blue arrow represents flow of paleodrainage.
60 significantly more robust data set for the deposition of the Punaschotter Conglomerates
within the Fiambalá Basin than existed previously, which placed onset of deposition
between 5.23 ± 0.44 Ma (Sample 055 in the Guanchin Formation along the Guanchin
transect, Carrapa et al., in press) and 3.77 ± 0.05 Ma (Sample 003 in the Punaschotter
Formation along the Guanchin transect, Carrapa et al., in press). These data also constrain
the age of the base of the Punaschotter more tightly than in any other basin along the
margin of the Puna Plateau.
2.6.2 Deposition, Folding, and Faulting in the Fiambalá Basin
Late Miocene-Pliocene units crop out along the western margin of the basin (Figures
9a-15a and 18). From north to south, the Punaschotter Conglomerates are well exposed along strike and were sourced from the north and east in the Rio Guanchin transect, and from the west, north, and east farther east in the basin (Carrapa et al., in press). The
Guanchin Formation is not exposed in the northernmost study section (Figures 9a & b), but crops out elsewhere along strike and also represents sources from the western, northern, and eastern margins of the basin, reflecting erosion of the Puna Plateau margin and western bounding ranges (Carrapa et al., in press). The Tamberia Formation does not crop out in the northern and central regions of the basin but is well-exposed in the south, with lithologic contributions from the western bounding ranges (Carrapa et al., in press).
Based on the unconformities and faulted and folded strata in the six constructed cross
sections (Figures 9b-14b), as well as cross sections presented in Carrapa et al. (2006; in
press), deposition of both the Punaschotter Conglomerates and its interbedded ash layers
occurred during extensive deformation of the Fiambalá Basin. From north to south, the
61
Figure 18. Geologic and structural mapping of the Fiambalá Basin. This figure is a compilation of all maps created for this study and encompasses figures 9-14.
62 amount of folding and faulting, and therefore the amount of shortening of late Miocene-
Pliocene units along the western margin of the Fiambalá Basin, increases. Cross sections
A-A’- C-C’ (north section, Figures 9b-11b) show gentle to moderate folding of all late
Miocene-Pliocene units, shaped into anticlines and synclines, and faulting either within the underlying basement unit, or to the west, where Paleozoic-Mesozoic units are thrust over Neogene units.
Moving southward, cross sections D-D’ and E-E’ (Figures 13b-12b, respectively) from the central section of the basin depict increased folding and faulting of all late
Miocene-Pliocene units, and Permian rocks begin to crop out to the west. The Tamberia
Formation is brought close to the surface by deformation and, in cross section E-E’
(Figure 12b), a west-vergent fault splay reaches the surface, thrusting the Guanchin
Formation over the younger Punaschotter Formation. This fault splay is not identified to the north, in cross section D-D’ (Figure 13b), and likely dies out before reaching this area. The fault likely dies out to the south as well.
Farther south, the Tamberia Formation is brought to the surface by deformation in cross section F-F’ (Figure 14b), and Permian rocks are thrust directly over Neogene strata, and Precambrian-Cambrian bedrock is present as well. Here, the Tamberia
Formation is thrust over the Punaschotter Formation by a steep east-vergent fault. The
Guanchin Formation has been severely eroded in the east, and the Punaschotter
Conglomerates unconformably overlie the Tamberia Formation in the center of the transect. To the west, the Guanchin Formation crops out minimally, and a dissected
Quaternary unit is observed blanketing the Tamberia Formation.
63 Farthest south, along the Rio Guanchin, the most intense deformation is observed, as
shown by Carrapa et al. (in press) (Figure 15b). Here, numerous east-vergent thrust faults disrupt the transect and carry the Tamberia Formation over the Guanchin Formation, and the Guanchin Formation over the Punaschotter Formation in some places. The Tamberia and Guanchin Formation beds are highly folded and, in places, the Tamberia Formation is overturned. The Punaschotter Formation has undergone moderate folding, though in places (notably where FA-9 and additional ash samples were collected) the Punaschotter overlies the Guanchin Formation at a nearly-horizontal orientation. Elsewhere, the
Punaschotter unconformably overlies the highly faulted and folded Guanchin and
Tamberia Formations.
These cross sections, particularly in the south, show that the underlying Tamberia and
Guanchin units have undergone higher degrees of deformation than has the Punaschotter
Formation. Greater deformation in the southern section of the basin than in the northern section has yet to be explained, but could be due to displacement profiles along faults where displacement tends to die out toward fault tips. Further study is necessary to quantify this.
2.6.3 Geochronology of the Punaschotter Conglomerates along the Puna Plateau
Margin
The Punaschotter Conglomerate or its equivalent has been found and dated in
numerous basins along the southeast-eastern margin of the Puna Plateau (Figure 19). In
the Hualfín-Río Basin to the west of Fiambalá, Punaschotter Conglomerates unconformably overlie both lower Neogene strata and Precambrian crystalline basement.
64
Figure 19. Constraints on the age of the base of the Punaschotter Conglomerates in regional basins of the Puna Plateau. Colors illustrate oldest (red) to youngest (light green) deposition of the conglomerates. 1. Fiambala Basin-western margin , 4.02 ± 0.04 Ma, U-Pb dating, (7 samples - this study). 2. Fiambala Basin- Rio Guanchin transect, 3.77 ± 0.05 Ma to 5.23 ± 0.30 Ma , U-Pb dating (4 samples - Carrapa et al., in press). 3. Hualfin-Rio Basin, 2.35 Ma (uncertain), based on calculation of constant sedimentation rate, 65 (Allmendinger, 1986). 4. El Cajon Basin, between 5.0 ± 0.42 - 6.93 ± 0.15 Ma, U-Pb dating (2 samples - Schoenbohm, in preparation). 5. Santa Maria Basin, 2.51 ± 0.06 - 2.96 ± 0.57 Ma, ZFT, (Strecker, 1989). 6. Angastaco Basin, 2.4 Ma, (Strecker, unpublished, in Coutand et al., 2006). 7. Quebrado del Toro, <8 - 4.17 ± 0.03 Ma, radiometric dating, (Strecker, 1997, unpublished, in Marret and Strecker, 2000). This boulder conglomerate is morphologically and compositionally identical to that in the
Fiambalá Basin and is interbedded with medium to coarse grained sandstone
(Allmendinger, 1986). The Punaschotter has not been directly dated here, but the underlying strata have. Marshall et al. (1981) and Butler et al. (1984) dated the Pliocene strata at ~3.54 Ma (40K-40Ar dating of a tuff sample, uncertainty unknown, Marshall et
al., 1981), which was later recalculated by Butler et al. (1984) to 3.53 ± 0.04 Ma at the
Puerta de Corral Quemado (40K-40Ar dating of a tuff sample). Using a constant
sedimentation rate of 56 cm/1000 years (Butler et al., 1984), Allmendinger (1986)
calculated that the age of the base of the Punaschotter Conglomerates, ~660 m above the
highest dated ash, at approximately 2.35 Ma.
The El Cajon-Campo del Arenal basins contain the Tortoral Formation, an alluvial
fan conglomerate that unconformably overlies the finer-grained Playa del Zorro
Formation. Recent U-Pb dating constrains the base of the conglomerate to 5.0 ± 0.42 Ma
- 6.93 ± 0.15 Ma (Schoenbohm, in preparation). The formation thickens westward and toward the center of the Chango Real Range and its bounding reverse fault, and is records
local alluvial fan sources (Mortimer et al., 2006). Eastward thinning is the likely result of
the uplift of the Sierra de Quilmes anticline, which prevented sediment from being
transported across the basin (Mortimer et al., 2006).
The Santa María Basin contains the Yasyamayo Formation, a conglomeratic unit that
unconformably overlies the Corral Quemado Formation. This conglomerate contains
clasts recycled from underlying strata (Bossi et al, 1984), and represents an areally
restricted alluvial-fan system shed from a nearby paleovalley to the northeast, Amaicha
66 (Kleinert and Strecker, 2001). Zircon fission track age constraints on ashes place the base of the unit at 2.51 ± 0.06 - 2.96 ± 0.57 Ma (Strecker et al., 1989).
The Angastaco Basin contains a conglomeratic gravel unit which unconformably overlies the conglomeratic San Felipe Formation, an upward-fining alluvial deposit
(Coutand et al., 2006). This conglomeratic unit, which contains recycled pyroclastic material at its base (Coutand et al., 2006) is thought to be the Punaschotter equivalent but has not yet been differentiated from the San Felipe Formation. Current work is being done to further evaluate the conglomerate. Dating of the pyroclastic material at the base of the unit by Strecker (unpublished, in Coutand et al., 2006) yields an age of 2.4 Ma
(uncertain).
The Alfarcito Conglomerate of the Quebrada del Toro Basin is a gray-colored, medium- to coarse-grained conglomerate intercalated with red-colored siltstone (Marrett and Strecker, 2000). The conglomerate conformably overlies Plio-Pleistocene fossil- bearing sandstones and intercalated Sola Conglomerates (Hilley and Strecker, 2005). A tuff within the Alfarcito Conglomerate has been radiometrically dated at 4.17 ± 0.03 Ma
(Strecker, unpublished data, 1997, in Marrett and Strecker, 2000). However, the conglomerate has been divided into two members, the lower and upper Alfarcito, due to differing flow directions during deposition. The lower Alfarcito is described as a clast- supported medium-bedded conglomerate with sub-rounded to rounded, well-imbricated clasts which lack a well-defined and repeating fining upwards sequence (Hilley and
Strecker 2005). This unit is dated at <8-4.17 Ma with a source lithology from the western
Puna margin. The upper Alfarcito is described by Hilley and Strecker (2005) as a clast- supported medium-bedded conglomerate with sub-rounded to sub-angular, well-
67 imbricated clasts and sandy intercalated beds. This upper unit has been dated at 4.17-0.98
Ma, with a source lithology from the more abundant and proximal Puncoviscana
Formation (Hilley and Strecker, 2005). Based on lithologic descriptions, the lower
Alfarcito appears to more closely resemble the Punaschotter Conglomerates of the
Fiambalá Basin than does the upper Alfarcito, although the essential differences appear to
be the more angular clasts and the presence of sandy intercalated beds in the upper unit.
The Punaschotter Conglomerates are described as well-rounded to sub-rounded and lack interbedded sandy units (Carrapa et al., in press). If this difference alone is sufficient to determine which portion of the Alfarcito is equivalent to the Punaschotter
Conglomerates, then the lower Alfarcito, dated between <8-4.17 Ma, represents the equivalent conglomeratic unit. If it is not, then the Alfarcito Conglomerate records an interval from <8-0.98 Ma, and requires further age constraints.
2.6.4 Comparison to Regional Basins
The Punaschotter Conglomerates in the Fiambalá Basin dated in this study are older
than lithologically and stratigraphically equivalent conglomerates in the Hualfín-Río,
Santa María, and Angastaco basins, yet younger than equivalent conglomerates in the
Toro and El Cajon basins (Figure 18). Previous authors have speculated on the cause of
conglomerate deposition in each of these regional basins; these interpretations are
presented here. Allmendinger (1986) suggests that deposition of the Punaschotter
Conglomerates in the Hualfín-Río Basin likely marked the culmination of uplift (~2 Ma)
in the immediate region because the unit unconformably covers older faults and folds.
Allmendinger (1986) speculates that the age of the conglomerate is probably not the same
68 everywhere (within the basin) due to morphological differences, but indicates that
deposition was the result of localized uplift.
In the El Cajon-Campo del Arenal basins, Mortimer et al. (2006) find that
conglomerate deposition coincides with basin dissection by the Sierra de Quilmes anticline, high-angle reverse faults, and progressive shortening from north to south. The
Tortoral Formation marks a facies change from interbedded sandstones and conglomerates to boulder conglomerates, which was due to increased sediment supply concurrent with or resulting from a climate change.
The Santa María Group of the Santa María Basin encompasses the Yasyamayo
Formation, and paleocurrent and provenance data suggest that the conglomerate was
transported from the Calchaquíes and Aconquija ranges to the east-northeast (Bossi et al.,
1984, 1993; Villanueva Garcia and Overjero, 1998). Uplift of these eastern ranges drove
folding and overthrusting of Neogene sediments after 3 Ma, and Kleinert and Strecker
(2001) suggest that the Yasyamayo Conglomerates were most likely deposited as the
result of tectonic uplift of adjacent ranges, and not regional climate change.
In the Angastaco Basin, late Pliocene shortening at 3.4-2.4 Ma due to uplift of the
eastern bounding Sierra de los Colorados range (Coutand et al., 2006) likely was responsible for changes in depositional environment (Carrapa and Mortimer, in preparation), leading to deposition of the coarse conglomerate. Carrapa and Mortimer (in preparation) suggest that such changes in depositional environment and increased aridity were due to tectonic deformation and basin reorganization.
In the Toro Basin, Hilley and Strecker (2005) found evidence that different source
lithologies of the lower and upper Alfarcito Conglomerates suggest changes in paleoflow
69 direction. Paleoflow changes were due to deformation within the Toro Basin, as well as uplift of the Sierra Pasha to the east, brought about by Miocene contraction and deformation (Marrett and Strecker, 2000).
2.6.5 Migration of a tectonic orogen
While the migration of the tectonic orogen from west to east across the Central Andes
(Figure 6) generally has been accepted as the driving mechanism creating uplift, exhumation, and formation of orographic barriers in the region (Coutand et al., 2006;
Deeken et al., 2006; Carrapa et al., 2005, 2006), it is unlikely that such migration is the cause of Punaschotter deposition along the Puna margin. The oldest ages of regional uplift (Figure 6) are in western to mid-western ranges, where uplift began in the Eocene-
Oligocene. Oligocene to Miocene uplift does appear to be somewhat out-of-sequence with uplift of surrounding ranges. Mid to late Miocene uplift occurred in the northeast, followed by late Miocene uplift in the southeast. If migration of the tectonic orogen was the cause of Punaschotter deposition, then deposition among basins should follow this same pattern (Figure 19). Instead, depositional ages within the Toro and El Cajon basins, both in areas of more recent uplift, are the two oldest ages (<8-4.17 Ma and 5.0 ± 0.42
Ma - 6.93 ± 0.15 Ma, respectively). Thus, deposition of the Punaschotter Conglomerates is not likely due to the eastward migration of the Andean tectonic orogen.
70 2.7 COMPARISON TO CLIMATE AND TECTONICS
2.7.1 The Fiambalá Basin Punaschotter Conglomerates
The onset of deposition of the Punaschotter Conglomerates in the Fiambalá Basin at
4.02 ± 0.04 Ma is on the older end of global climate change and associated increased
sedimentation and erosion rates, and increased clastic grain size at ~2-4 Ma, recognized
by Molnar and England (1990), Peizhen et al. (2001), and Molnar (2001). Regional
climate history demonstrates increased aridification in western South America at ~6.4-2.6
Ma (e.g. Hartley and Chong, 2002; Lamb and Davis, 2003; Gaupp et al., 1999), due to
changes in the Hadley circulation, increased upwelling of the Humboldt Current, and
establishment of the Panama land bridge after 3-3.5 Ma. Along the eastern Puna Plateau
margin, a transition from humidity to aridity was recorded at 3.4-2.4 Ma (Coutand et al.,
2006; Kleinert and Strecker, 2001) in the Angastaco and Santa María basins, though a
transition from arid to hyper-arid conditions was also recorded in the late Miocene-early
Pliocene in the Angastaco and Fiambalá basins (Carrapa and Mortimer, in preparation;
Carrapa et al., in press). These local and regional differences in climate illustrate the fact
that climate can have differing effects in different locations. However, deposition of the
Punaschotter Conglomerates in the Fiambalá Basin at 4.02 Ma generally aligns with
global, regional, and local evidence of climate change and increased aridification. Not surprisingly, deposition most closely aligns with regional and local climate changes,
because climate changes in such close proximity to the study site should have a greater
effect on sediment deposition than should climate changes on a global scale.
Fiambalá Basin bounding ranges were inferred to have uplifted during or by the late
Miocene (~7-6 Ma, AFT), though deposition of the Punaschotter Conglomerates did not
71 begin until ~4.02 Ma. If uplift of basin-bounding ranges was the driver of Punaschotter
deposition, then deposition should have commenced at the time of local uplift. Because it
did not, climate almost surely had to exert a role on deposition. We propose the following
explanation: as basin-bounding ranges began to be uplifted in the late Miocene,
particularly those ranges to the east, an orographic barrier was created that blocked
Atlantic-derived precipitation from reaching the basin. Over time, aridity increased across
the basin. By ~4 Ma, aridity was great enough to cause deposition of the Punaschotter
Conglomerates. Deposition of this unit was most probably the result of tectonic uplift
driving climate change and increased aridity. However, alignment between deposition of
the Punaschotter and both global and regional climate changes cannot be overlooked.
Both regional, and possibly global, climate changes likely enhanced aridity in the
Fiambalá Basin in the late Miocene-early Pliocene, and may have worked in concert with tectonically-driven aridity, resulting in deposition of the Punaschotter Conglomerates.
2.7.2 Regional Basin Conglomerates
If global climate change was the cause of increased erosion and sedimentation rates,
leading to conglomeratic deposition (increase in clastic grain size), then deposition of
similar conglomeratic units throughout the Puna Plateau margin should correlate as
robustly with climate change as conglomerate deposition in the Fiambalá Basin does. In
addition, conglomerate deposition should be synchronous among basins if deposition is
the result of regional or global climate changes. Deposition of the Punaschotter
Conglomerates in the Hualfín-Río, Santa María, and the Angastaco basins correlate with
global climate change at 2-4 Ma, although deposition occurred on the young end of that
72 climate change. However, deposition in the Toro and El Cajon basins does not correlate
as strongly and the ages of deposition within these two basins are marginally older, if not
significantly older, than global climate change. These ages may represent times of clastic
coarsening and increased deposition due to a regional transition into aridity at 6.4-2.6 Ma.
Ages of deposition in the Toro and El Cajon basins may also have been the product of
regional tectonic uplift, and deposition may have been independent of climate change.
Collectively, deposition of the Punaschotter Conglomerates across the regional basins
was not synchronous with global climate change from 2-4 Ma.
Regionally, deposition of conglomeratic units correlates moderately well with
increased aridity between 6.4-2.6 Ma. The Hualfín-Río, Angastaco, and Santa María
conglomerates were deposited near the very end of an episode of increased aridification, while conglomerates in the El Cajon and Toro basins were deposited at the very early stages of aridification, if not before. A transition from semi- to hyper- aridity was interpreted by Hartley and Chong (2002) and Alpers and Brimhall (1988) for the
Atacama Desert at ~11-7 Ma and ~14-9 Ma, respectively. Deposition in the El Cajon and
Toro basins may have been affected by earlier aridification, though it is unclear why such
regional climate change would have only affected these two basins and not the others.
Deposition in the Hualfín-Río, Angastaco, and Santa María basins did respond to climate
change until the most heightened period of aridification. If regional climate was the
driver of conglomerate deposition, it did not have the same effect in all basins along the
Puna Plateau margin, thus making it difficult to clearly link regional climate change to
deposition of this unit.
73 Deposition of the Punaschotter Conglomerates among basins aligns with regional climate change more strongly than it does with global climate change. This makes sense because regional climate should have a stronger effect on regional deposition than should global climate. Additionally, regional climate change extended over a longer period of time than did global climate change. Overall, a link between climate and deposition in some regional basins is strong, while in other basins, the link is much weaker. Such findings do not rule out a correlation to regional climate change entirely, yet render strong correlation between regional climate change and conglomerate deposition tentative in at least in two of the basins along the eastern margin.
If uplift of basin-bounding ranges was responsible for the deposition of the
Punaschotter Conglomerates along the plateau margin, then deposition should occur simultaneously with inferred uplift, or at least follow uplift closely. In many cases, we find that deposition did not occur simultaneously with local uplift, but rather near the end of local uplift or after local uplift ceased. Deposition in the Santa María basin did not begin until 2.9-2.5 Ma, well after uplift of bounding ranges began in the late Miocene
(~7-6 Ma). Deposition of the Punaschotter Conglomerates in the Hualfín-Río basins by at least 2.35 Ma likely marked the culmination of uplift of the Sierra de Hualfín around 2
Ma. Deposition in the Angastaco basin at 2.4 Ma occurred at the end of late Pliocene uplift of the eastern bounding Sierra de los Colorados (3.4-2.4), though well after onset of mid-to-late Miocene uplift of the Sierra Quilmes Anticline. In the El Toro Basin, the
Alfarcito Conglomerate was deposited between <8-4.17 Ma, and deposition may have begun near the onset of uplift of the eastern bounding Sierra Pasha range (8-6 Ma), may have followed uplift, or may have preceded it. Deposition lags behind uplift of the
74 western bounding Cumbres Luracatao, which began to be uplifted in the Oligocene-
Miocene. Deposition of the Tortoral Formation in the El Cajon basin after 5-6.93 Ma
occurred concurrently with local uplift in the late Miocene.
As in the case of the Fiambalá Basin, deposition in many of the basins did not begin
at the onset of bounding range uplift, but rather well into localized deformation. If the
same scenario applied to the Fiambalá Basin is applied to other regional basins, a pattern can be seen. In the Hualfín-Río, Santa María, and Angastaco basins, bounding range uplift commenced at >4 Ma, 9-4 Ma, and 13-3 Ma, respectively, before deposition of the conglomerates. The uplifts likely drove increased aridity by creating orographic barriers.
Global and regional climate changes more or less correlate to the deposition of conglomeratic units near the end of aridification and may have enhanced aridity throughout each of these three basins.
The El Cajon and Toro basins experienced deposition that may have been concurrent
with bounding range uplift, along one if not all margins of each basin. This may have
been due to differences in the effect of orographic barriers in these locations. In addition,
earlier response aridity may have already been in place in these two locations before
uplift of bounding ranges commenced, shortening the response time to deposition of the
conglomerates.
2.8 SYNTHESIS
Deposition of the Punaschotter Conglomerates in the Fiambalá Basin fits into an
episode of global climate change at 2-4 Ma, regional climate change from 6.4-2.6 Ma,
and deformation of Neogene units in the early Pliocene. Because deposition did not begin
75 at the onset of bounding range uplift along the Fiambalá Basin, uplift alone is not likely
to have been the sole driving force of conglomerate deposition, and must have been
coupled with another force, climate, to create increased clastic grain size. As local ranges were uplifted, orographic barriers were created, increasing local aridity. Regional and global climate change likely enhanced aridity and resulted in deposition of the
Punaschotter Conglomerates within the Fiambalá Basin.
Conglomerate deposition in other regional basins is asynchronous between these
basins and with conglomerate deposition in the Fiambalá Basin. Deposition in some
basins fits well into the global and/or regional climate change window, yet the fit is not
good in others. The onset of uplift of basin-bounding ranges significantly precedes
conglomerate deposition in most of the basins, but correlates well with conglomerate
deposition in others. Tectonics and climate acting independently cannot explain
deposition of the conglomerates in all regional basins. However, when coupled together,
tectonics and climate provide a more probable cause for deposition of these units. The
scenario suggested for deposition of the Punaschotter Conglomerates in the Fiambalá
Basin can be applied to deposition of conglomeratic units in other regional basins.
Our study finds that regional tectonics is the most likely force capable of causing
increased erosion, sedimentation, and clastic grain size, and is the most probable driving force behind deposition of the Punaschotter Conglomerates. As basin-bounding ranges were uplifted along the eastern margin of the Puna Plateau in the late Miocene-early
Pliocene, orographic barriers formed east of the regional basins and eventually blocked
Atlantic precipitation from many of the basins. The amount of precipitation received by each basin varied, depending on its location and proximity to orographic barriers, and the
76 existing climate regime. Timing of deposition also depended on the degree of aridity
already affecting each basin. Tectonically-driven aridity was likely enhanced most
strongly by regional climate-driven aridification at 6.4-2.6 Ma and, to a lesser degree, by
global cooling at 2-4 Ma. Deposition of the Punaschotter Conglomerates in the Fiambalá
Basin, as well as in other regional basins, was likely the product of coupled tectonically-
driven aridity, caused by the formation of orographic barriers, and increased aridity due to both regional and global climate cooling, resulting from a disequilibrium in the landscape, all driven by large-scale tectonic changes and variations in orbital parameters.
77
CHAPTER 3:
CONCLUSION
3.1 SUMMARY
This study has more tightly constrained the age of onset of deposition of the
Punaschotter Conglomerates in the Fiambalá Basin to 4.02 ± 0.04 Ma. Combined with previous ages by Carrapa et al. (in press), a robust data set now exists for the basin and has been compared to data from other regional basins to make a stronger assessment of the timing of regional deposition of this unit. U-Pb dating of seven interbedded ashes and ignimbrite samples, in combination with ashes dated by Carrapa et al. (in press), suggest that two possible eruptions deposited ash throughout the basin during deposition of the
Punaschotter Conglomerates. If this is the case, along-strike variations in the
Punaschotter-Guanchin contact may have been due to localized variations in tectonic uplift rate, short term events such as landslides, floods, earthquakes, or differences in paleowatershed size. Detailed mapping and cross sections reveal that the Punaschotter was deposited during extensive folding and faulting in the Fiambalá Basin, and that deformation was more severe in the southern section of the basin then in the north.
Deposition of the Punaschotter Conglomerates within the Fiambalá Basin correlates with episodes of global and regional climate change, at 2-4 Ma and 6.4-2.6 Ma, respectively. Deposition also occurred during deformation of Neogene units within the
78 basin, although deposition did not commence at the onset of bounding range uplift. If evaluated independently of other regional basins, a clear connection could be drawn
between global and regional climate change and deposition of the Punaschotter
Conglomerates in the Fiambalá Basin.
However, deposition of the Punaschotter Conglomerates was asynchronous among regional basins, and asynchronous with global climate change. For this reason, global climate change cannot have been the driver of conglomerate deposition along the Puna
Plateau margin. Deposition in regional basins correlates moderately well with regional
climate change, though some exceptions exist. We find that deposition followed the onset
of uplift of basin bounding ranges by a few millions of years, in most cases, and that
tectonics, coupled with climate change, is the most reasonable explanation for deposition
of the conglomerates.
We propose that regional tectonics were the most probable force driving deposition of
the Punaschotter Conglomerates. Late Miocene-early Pliocene uplift of basin bounding ranges along the Puna Plateau margin increased local aridity by forming orographic barriers, which blocked Atlantic-derived precipitation from reaching the basins. Over
time, aridity increased across the basins, resulting in deposition of the Punaschotter
Conglomerates or its local equivalents. While deposition originally resulted from tectonically-driven aridity, global and regional climate change, driven by continental scale tectonic forcing and global scale orbital forcing, may have enhanced aridity across the margin, thus enhancing deposition of this conglomeratic unit.
79 3.2 LIMITATIONS
There are topics that this study addresses only briefly, or not at all, and issues that
remain unresolved. Our discussion of uplift of large-scale landmasses, such as plateaus,
addresses the central issue of tectonics driving climate change. Other facets of this
process discussed briefly include changes in distance of incoming and outgoing solar
radiation due to landmass uplift, atmospheric pressure and circulation, climate models,
and associated changes in sea level pressure and precipitation rates. The evolution of
orographic barriers is cited in terms of topographic changes that drive changes in
atmospheric circulation and precipitation, ultimately leading to climate changes. Changes
in air temperature and moisture, windward and leeward air circulation, and rain shadow
effects are mentioned only briefly. A more-in depth discussion of orographic barriers and
plateau uplift can be found in Raymo and Ruddiman (1992), Hay (1996), Ruddiman
(1997), and Coutand et al. (2006).
The links between orographic barriers and rates of exhumation and precipitation, sediment erodability, erosional efficiency, and drainage reorganization are intricate and detailed, and only selected aspects of this issue were included in this discussion.
Discussions by Sobel and Strecker (2003) and Carrapa et al. (2005) offer greater insights
into these issues.
Feedback mechanisms relating to snow cover and albedo, cooling temperatures,
glaciations, Antarctic ice sheet expansion, and ocean and atmospheric circulations are all
given brief mention, though the details of each process are not discussed here. Detailed
discussions of all processes can be found in Mercer (1983), Raymo and Ruddiman
(1992), Hay (1996), Maslin et al. (1996), and Ruddiman (1997). Additionally, little
80 discourse is given to the effect of drifting continents, the opening and closing of oceanic
gateways, volcanism, and the introduction of CO2 and SO2 into the atmosphere. Hay
(1996) offers an in depth analysis of each.
Deposition and development of alluvial fans within the basin, and their tectonic and
climatic signals are not discussed, though a good discussion can be found in Quigley et
al. (2007). The effects of short-term geologic events such as earthquakes, landslides, and
floods on sedimentation and erosion are mentioned briefly. In order to more accurately
quantify the effects of these processes, additional studies and data collection are
necessary. Additional information on these occurrences in other regions of the world can
be found in Frostick and Reid (1989), Hovius et al. (1997, 2000), Keefer (1994), and
Molnar (2001).
3.3 DATA ISSUES
Despite our best efforts, there are issues that have not been addressed and potential
“gaps” in our analysis. No data was collected from the eastern margin of the basin, due
largely to time constraints and a lack of Neogene outcrops. At sample locations FA-8 and
FA-9, multiple layers of ash were found, yet only the uppermost layers were sampled.
Our study did not extend farther south than the Rio Guanchin, leaving the southernmost
section of the western margin for further investigation. Regions of the western margin exist that we did not investigate, due to time and transportation constraints. Additional samples from “gaps” between samples sites could better constrain the Punaschotter-
Guanchin contact variation. An in-depth watershed analysis using computer modeling
may also aid in correlation between age of conglomerate deposition and paleowatershed
81 size. During sample analysis, lab error is always a potential, particularly in dealing with zircon grains on micron scales. Analysis points of each zircon were hand-selected, allowing for possible error if selection points were too close to the center of the grain.
However, the use of zircon calibration standards, weighted average calculations, and
other error-reducing techniques lowers the possibility for human error. Sample FA-4 only
had 4 viable zircon grains for analysis, reducing the accuracy of the age. Additionally,
little focus was given to zircon crystallization rates.
3.4 Future Work
Suggestions for further study in the Fiambalá Basin include ash collection along the
eastern margin of the basin, further collection of ash samples in locations with multiple
layers, and investigation into the southernmost region of the basin, as well as transects
not reached in our study. Analysis of the source volcano and eruption of the ash layers, in
addition to more robust age constraints of samples with fewer than ten grains, would
strengthen the data. Seismic reflection data does not exist in the Fiambalá Basin, as it
does in the El Cajon-Campo del Arenal Basins. Such data would improve subsurface
interpretations and may better constrain basin history. Investigations on the effects of
earthquakes, landslides, and floods would be useful in more accurately analyzing
sedimentation rates and overall basin history. Further analysis of watershed size would
be useful in addressing the issue of age variation in the Punaschotter-Guanchin contact,
as would a better understanding of the driving mechanisms behind the greater deformation of Neogene beds in the southern basin than in the north.
82
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