Download Text S1 S1. Pampa del Tamarugal 20°

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S1. Pampa del Tamarugal 20°-22°S
S1.1. Stratigraphy
For the region between 20°10’S and 21°50’S there are few published maps and reports.
This is where our research has focused. Although the Oligocene through Middle Miocene rocks
crop out only locally, mostly in the canyons that cut the Western Andean Slope, their relative
positions and thicknesses have been determined from boreholes and their spatial extent is known
from seismic mapping [Victor et al., 2004; Nester, 2008]. Upper Miocene and Pliocene strata
and unconsolidated sediments are much more widespread on the surface of the basin (Fig. 7),
though thin.
The available ages of the Altos de Pica Formation include three in the upper member
(AdP member 5) (Table S1). Two ash interbeds within ~15 m of the local top of AdP member 5
(Table S1, samples IB24 and IB34, located ~55 km distant from one another) yield dates of 15.8
± 0.08 Ma and 12.9 ± 0.08 Ma (ages are reported with 2-sigma analytical errors). Though
another tuff (Table S1, IB20) <50 m below the top of the AdP member 5 yielded Ar/Ar data that
suggest a somewhat younger age, those data are difficult to interpret, varying from a total gas
“age” of 11.62 ± 0.1 Ma to a pseudo-plateau of approximately 13 Ma. Consequently, pending
new analytical results, we interpret an age no younger than 12 to 11 Ma for the end of deposition
of AdP member 5 in this region. In the text, we simplify this to the statement that the top of AdP
is as young as ~11 Ma.
Poorly consolidated Upper Miocene alluvial, salar and, locally, eolian deposits overlie the
AdP. Widely exposed in the eastern half of the lowland region, west of the Precordillera, are
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uppermost Miocene to lowermost Pliocene alluvial deposits (Fig. 7). At 21°40’S they comprise
the Arcas Fan, whose age Kiefer et al. [1997] considered to span ~7.2-6.8 Ma based on K-Ar
ages in a pair of associated volcanic deposits (Table S1). Matrix-supported and poorly sorted
conglomerates dominate the Arcas fan strata [Kiefer et al., 1997] and characterize the strata
associated with the oldest relict alluvial landforms elsewhere in the valley [Nester, 2008]. We
refer informally to this stratigraphic unit as the Arcas unit. At elevations between about 2300 m
and 1900 m, these strata occur within the major canyons, filling broad paleochannels eroded into
the older Miocene strata, whereas below about 1900 m they spread widely across the lowlands.
Although thicknesses of only ~10-50 m of Arcas unit are exposed in canyon walls, subsurface
data [e.g., Kiefer et al.’s 1997 analysis of gravity data; Nester’s 2008 analysis of seismic
reflection profiles] indicate that 300-600 m thickness of Arcas unit and younger strata occurs
below the surface of the valley. A pyroclastic deposit within the Arcas unit, the Carcote
Ignimbrite, has been dated with Ar-Ar methods on biotites, yielding ages of 5.38 ± 0.09 Ma and
5.26± 0.04 Ma [Hoke et al., 2007; Nester, 2008, respectively] (Table S1, samples IB-09 and IB17, respectively). Atop this ash at the latitudes of this study are typically another 20-30 meters of
alluvial deposits. Based on the age of the Carcote Ignimbrite, we assign an age of ~5 Ma for the
cessation of deposition for this last large-volume phase of latest Miocene-earliest Pliocene
deposition. The Arcas deposits are intercalated in the western sector of the basin with the lower
units of a sequence of Upper Miocene – Lower Pliocene(?) lacustrine limestone and salt pan
evaporite deposits [Sáez et al., 1999]. After the deposition of the Arcas unit, the primary
aggradational phase within the basin ceased. Nevertheless, thin veneers of Pliocene and
Quaternary alluvial fan deposits occur west of the major canyons and grade into salt pan and
lacustrine deposits near the western limit of the valley.
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Eolian sandstone is common in the northeastern extreme of the study area (Fig. 7), draped
across the western flank of the Precordillera. Between 20°28’S and ~21°S eolian facies
sandstone dominates an outcrop belt 7 to 13 km wide (east-west direction) at elevations up to
2900 m. The eolian sandstones overlie the AdP across a surface of very low angle erosional
truncation. Hundred-meter-thick climbing translatent stratification in the eolian beds creates the
false impression that the AdP-eolian contact is sharply angular. Individual sets are tens of meters
thick. The eolian sandstone reaches 100-150 m thickness near its eastern limit, but thins to only
50 m thickness over short distances to the east and west, defining the NNW-trending crest line
and flanks of a major dune field. Fluvial and alluvial conglomerate are associated with the eolian
deposits, most likely comprising two stratigraphic units. To the east of the eolian axis, light gray
weathering conglomerates stretch from the line of dunes across the AdP top; the easternmost
patches occur at ~3900 m altitude. This conglomerate is 50-100 m thick over a broad region, and
displays a mixture of channelized, sheet flood, and debris flow fabrics. Overlying and to the west
of the dune deposits is a dark gray-weathering conglomerate, also ~100 m thick, which displays
sheet flood and debris flow beds. This conglomerate extends to elevations below 1900 m, where
it is overlain by gravels that constitute relict alluvial fan landforms, which we assign to the Arcas
unit. Pending more extensive study (in progress, N. Blanco), we interpret that growth of a dune
field created a topographically isolated valley to its east in which the most proximal sediments
were trapped by a locally high baselevel. In that eastern valley, a proximal gravel accumulated
(the high altitude light gray conglomerate) that interfingered westward with eolian sandstone.
The western, dark gray conglomerate unit filled the topographic low on the western side of the
dune field and eventually overtopped the dunes, but we do not know yet whether this western
gravel was entirely younger than the dune field or partly contemporaneous. Two interbedded
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volcanic ashes in this complex of units have been dated thus far, yielding an age of 8.5 +/- 0.5
Ma in the eolian deposits, and about 5.6 +/- 0.2 Ma in conglomerates overlying the paleo-dune
deposits (Table S1, Supplemental Data) (Blanco and Tomlinson, work in progress). We consider
these strata to be dominantly of Late Miocene age, though in acknowledgement of the fact that
we do not have a constraint on the minimum age we map the eolian sandstone and gravel unit as
a “Miocene-Pliocene” unit (Figs. 7, 8).
These chronological data and unconformities reveal that the Late Miocene-Pliocene
sedimentation in the eastern Pampa del Tamarugal basin occurred during three (?) discrete time
intervals. If the end of AdP deposition is as young as ~11 Ma (rather than ~12 Ma), then the
earliest interval of Late Miocene activity is captured in the terminal levels of the Altos de Pica
Formation. The second interval was a localized phenomenon caused by the creation of a major
dune field in the northeast that created a localized sediment trap to its east. The Arcas unit
accumulated during a third time interval, spanning the Miocene-Pliocene limit. Although the
numerical dates for the Arcas and for the eolian unit and its neighboring conglomerates do not
yet resolve a clear difference, the Arcas sits in a distinctive landscape position and appears to
overlie the eolian-conglomerate unit above an erosional unconformity. South of the area with the
eolian sandstone and associated conglomerates, an erosional unconformity separates the AdP5
and Arcas unit. In the absence of specific data to constrain the age of the oldest strata in the
Arcas unit, we interpret that several million years elapsed between the end of AdP5 and the
initiation of the Arcas unit.
This succession of stages of landscape activity in the southern Pampa del Tamarugal is
similar to the stages reported by Evenstar et al. [2009] in the northern Pampa del Tamarugal.
Most likely the differences between our numerical ages of landscape stages compared to their
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ages results from the difference between our dating method (interbedded volcanic deposits in
stratigraphic units) and that of Evenstar et al. [2009] (cosmogenic 3He dating of clasts on the
surface).
S1.2. Data and Measurements
Fourteen paper seismic lines for the southern Pampa del Tamarugal originally obtained
by Evergreen Resources and ENAP were provided by International Exploration Associates Ltd..
Stacking velocities were noted on the seismic lines. The paper seismic lines were scanned,
converted to bitmap, and eventually converted to segy format using the MATLAB programming
language, the Seismic Unix program, and ProMax version 2003.3.3, and then imported into the
seismic interpretation software Kingdom Suite version 8.1 (for a detailed workflow, see Nester
[2008], Appendix A.) Once in Kingdom Suite, reflectors were tied to available well log data,
located at the western edge of the seismic grid in the western portion of the basin. Tested at the
two wells, the stratigraphic matches were of moderate quality [Nester, 2008].
The modern dips of ancient landscape surfaces were calculated using depth profiles of the
paleosurfaces associated with the end of AdP member 5 deposition and the end of the Late
Miocene depositional episode (Figure 8). We quantify the form of those paleosurfaces using both
surface data and seismic reflection data. For the western subsurface part, we constructed profiles
for the four east-trending seismic lines that do not cross surface-breaking faults with significant
offset near the toe of the Western Andean Slope and along which the AdP5 surface appears
smooth and continuous over a distance of 10’s of kilometers (seismic lines 99-06, 99-07, 99-09
and 99-10, see Fig. 7). Seismic travel time to depth conversion using interval velocities
(calculated from stacking velocities) at every 100 shotpoints (2.5 km horizontal distance)
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resulted in depths to the reflectors. For the eastern surface exposures, geologic profiles were
constructed to extend east from these seismic lines to the up-dip limit of the surface of the AdP
member 5 deposits. Each surface section was constructed for a pair of parallel topographic
profiles, spaced less than 1.5 km apart, one located to cross the major canyons and its pair to lie
largely on the pediplain between the canyons. We utilized our geologic mapping results near the
surface profiles (field observations and remote sensing interpretations) to place stratigraphic
units in the topographic profiles (Figure 8). We used our knowledge from field mapping near the
traces of the seismic lines to tie surface stratigraphy to reflections, especially the locations where
reflectors intersect with the land surface. For the four profiles, the surface geologic controls
account for ~30% to 75% of the full width of the region between the eastern limit of the AdP and
the western monoclinal hinge.
Comparison of the paleosurface corresponding to the top of the AdP in the subsurface
and in outcrop reveals that their forms are very similar, but that the apparent dips are not
identical (Fig. 8). For three of the four profiles, there is an apparent increase in dip in the seismic
data compared to the surface data (Fig. 8a,b,c), and for line 99-10 (Fig. 8d) where there is spatial
overlap between the seismic data and canyon-wall exposures, there is a mismatch of the depth of
the easternmost subsurface control point for the top of the AdP relative to its exposed position.
These observations reveal the difficulty of quantifying the position in depth, and consequently
the inclination of the subsurface paleosurfaces, because of the significant uncertainty on the
seismic velocities approximated from stacking velocities. Consequently, we quantify the rotation
history based on the inclinations of the broad expanses of the paleosurfaces where they are
exposed at the surface and mapped using field and remote sensing observations.
To establish the position of the western hinge of the monocline, of most importance is the
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change in depth between successive points on the seismic profile, which define the inclination of
reflectors of interest. Consequently, concerns about the magnitude of the error on the seismic
velocity focus on whether the errors of stacking velocity conversion are consistent within a given
seismic line.
The modern dip of the ancient landscape corresponding to the end of deposition of the
Upper Miocene-Pliocene Arcas unit is extracted from the surface topography of the pediplain. As
pointed out by Hoke et al. [2007] and Evenstar et al. [2009], the pediplain is polyphase. Our lines
of topographic profile (Fig. 7, 8) transect landscapes formed in the early(?) Late Miocene, early
Pliocene, and Pliocene(?). However, in the sector from which we extract the inclination of the
Arcas unit across the limb of the large-scale monocline, the profiles transect either only Arcas
unit (profiles 99-09, 99-10) or Arcas unit and Upper Miocene-Pliocene eolian sandstone and
gravel (profiles 99-06, 99-07).
Summarized on a map, the lower or western monoclinal hinge occurs throughout the
study area where the modern elevation is between approximately 1500 – 1750 m (Fig. 7).
Potential tilting of key horizons caused by differential burial compaction of the strata was
investigated using thickness data (z) from the depth-converted seismic lines (99-07, 99-09 and
99-10). The Cenozoic strata subject to compaction reach local thickness maxima as great as ~950
and ~1200 m (based on the stacking velocity conversions) [Nester, 2008]. Nevertheless, during
the ≥11 million years for which our analysis documents rotations, the relatively thin Arcas unit
added little mass to the top of the AdP and, therefore, compaction would have been minor after
~11 Ma. This is because the Arcas reaches only 320 m thickness locally across the lower sector
of the monocline where the seismic lines image the geometry. For example, using the
OSXBackstrip software [N. Cardozo, 2007, freeware] to estimate the total compaction (assuming
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a depositional porosity of 35%, and a “c” value of 0.4), the AdP strata beneath the ~11 Ma
surface would have compacted at the western hinge of the monocline by only ~31 m, accounting
for a compaction rotation of only 0.2° ± 0.1° (means and standard deviation of values for four
seismic lines considered). We treat (Table 1) the compaction tilt uncertainty to equal this value,
0.2°, to reflect the range of possibilities between there having been no compaction and this
reasonable but untestable value from our numerical experiments.
2. Salar de Atacama Basin
S2.1. Stratigraphy
The age of the base of unit L is not well constrained. At the northwestern extreme of the
Salar de Atacama basin, 17 Ma volcanic rocks overlie folded rocks of Oligocene to earliest
Miocene age [Ramírez; 1979; Mpodozis et al., 2000]. By analogy to that angular unconformity,
the gentle angular unconformity at the base of L in the northeastern part of the basin may be of
Early Miocene age. The large uncertainty on this estimate is somewhat diminished by Hartley
and Evenstar [2009]’s demonstration that a late Early Miocene angular unconformity is
widespread in northern Chile. Reutter et al. [2006] interpret the age of the base of L very
differently than Jordan et al. (2007). Reutter et al. [2006] identify an unconformity beneath the
eastern part of the basin which they interpret to separate Eocene and older strata from Oligocene
and younger strata (their Fig. 14.4). In contrast, Jordan et al. [2007] identify the same
unconformity as the contact separating the Oligocene and Lower Miocene strata (their lower unit
K) from Middle Miocene and younger strata (our unit L). Were the Reutter et al. [2006]
interpretation correct, the age of long-wavelength rotation and tectonic uplift documented here
would span the Oligocene as well as the Miocene.
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All the sedimentary rocks from the salar surface to a base chosen where there is a major
lithologic change in the Toconao-1 borehole constitute unit M, which is dominated by halite.
Unit M reaches over 1800 m thick in the northern extreme of the salar and over 1400 m thick in a
second depocenter in the southern half of the salar (Fig. 10).
Ignimbrite interbeds near the eastern limit of units M and L are documented locally in
boreholes and, over a broader area, inferred from seismic reflection characteristics [Jordan et al.,
2002]. The occurrence of interbedded ignimbrites suggests ages younger than 10 Ma [de Silva
1989], but the maximum plausible age is uncertain because surface geology does not clarify the
age of the oldest major ignimbrites in the region. Ignimbrites as old as approximately 10 Ma are
known [Lindsay et al., 2001], but insofar as all older units east of the salar are buried, there may
exist buried older ignimbrites. We will refer to the base of M as no older than early Late
Miocene (~10 Ma) in age [Jordan et al., 2002].
S2.2. Tilted Long-Wavelength Surfaces and Relief Growth
Of 21 seismic lines that are oriented approximately east-west (N92°E to N101°E) and
that traverse the eastern half of the salar, we have had access to 11 of these, spaced along the full
north-south distance of the salar. Among those located north of the area of basement uplifts
(north of 23°28’S), the primary structure of the basin is of long-wavelength westward-fanning of
strata. However, the majority of the seismic lines display either short-wavelength folds and faults
related to the Peine fault system, or poor quality reflections at their eastern extreme where
seismic fold decreases and facies change from salar facies on the west to alluvial fans and
volcanic units on the east. Both conditions interfere with our ability to trace the bases of M and L
and to measure long-wavelength rotation. Across short-wavelength structures the inclination of
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units M and L are a sum of local and long-wavelength conditions, and we have insufficient data
with which to separate those two signals. As a consequence, we are relatively sure that we can
differentiate and quantify the long-wavelength tilt for only 3 of the available lines (Table 2;
Tables S2, S3), for the line segments a, b, and c that are marked on Figure 10.
S2.3. Data and Estimation of Uncertainties
Conversion of seismic profiles to the depth domain: To convert Two Way Travel Time
(TWT) to depth, we have only the sonic velocity logs from the Toconao-1 borehole in a central
part of the salar as data [Jordan et al., 2007]. To demonstrate the degree to which variability of
velocities in the eastern part of the salar might influence the measured tilt, we tested two velocity
models that bracket the model based on the Toconao-1 well data (Table S4). Velocity model #2
considers the case of higher velocity rocks, like may be found if the drilled salar-center facies
interfinger eastward with volcanic rocks. Velocity model #3 considers lower velocity rocks, as is
anticipated if the halite of the Toconao-1 borehole interfingers eastward or northward with more
siliciclastic or gypsum-rich facies. For all three models, we calculated the dip of the bases of M
and L assuming constant velocities between the western monoclinal hinge and the eastern control
point on the seismic profiles. Relative to the Toconao-1 velocity model (mean relief growth
based on unit M of 2300 ± 770 m), the higher velocity model causes an increased estimate of
long-wavelength relief growth (2590 ± 870 m), and lower velocities suggest less monoclinal
relief growth (1480 ± 540 m), giving a mean of ~2100 m and an uncertainty of ~1280 m (Table
S3). However, the true sonic velocity variations of units M and L in the Salar de Atacama basin
are likely much more complex spatially, both from center line to eastern margin, and from the
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southern part of the salar to the north, than has been probed here. The uncertainty on the sonic
velocity generates a large uncertainty on the magnitude of relief growth.
Compaction rotation: Nonmarine sediments deposited in subaerial environments
accumulate with lower initial porosity than is true of marine sediments of equal grain size
[Nadon and Issler, 1997]; all but about 100 m of the 3700 m of strata below unit L transected by
the Toconao-1 exploration borehole are nonmarine [Muñoz et al., 2002; Jordan et al., 2007].
Cenozoic units exposed along the western rim of the Salar de Atacama basin tend to be complex
alluvial facies [Pananont et al., 2004; Mpodozis et al., 2005], which typically would have had
depositional porosities significantly less than that of a well-sorted, well-rounded sandstone
(typically 40-45%) [Damanti and Jordan, 1989; Nadon and Issler, 1997]. Flint (1987) documents
the diagenesis of unit K sandstones, showing that vigorous diagenesis during burial reduced
porosity to less than 5%, and that 5% to 30% of the volume consists of cements. The latter
observation implies that porosity loss was only partially caused by compaction, but the
variability of cement percent implies that much of the porosity reduction would have led to
subsidence of overlying horizons.
We exclude Cretaceous rocks from the burial compaction estimation because they had
been strongly folded and erosionally beveled prior to Cenozoic accumulation, a history that
likely resulted in an advanced state of irreversible compaction. While diagenesis would have
continued to change the cement phases and permeability of the Cretaceous strata throughout the
Cenozoic, changes in volume would no longer have been coupled in a predictable manner to
increased burial depth.
We quantified the uncertainty on the amount of tilt caused by compaction by
decompacting the strata using a family of possible compaction curves. We used the freeware
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OSXBackStrip [Nestor Cardozo, 2005,
http://www.homepage.mac.com/nfcd/work/programs.html], which computes compaction under
burial load for user-specified choices of depositional porosity and of curvature of an exponential
porosity-depth curve [Sclater and Christie, 1980]. Based on Flint [1987] and Nadon and Issler
[1997], we used a depositional porosity of 35% and varied the constant, c, that controls the shape
of the porosity-depth function between values mimicking no compaction (c=.0001) and values
that cause compaction to only 5% inter-grain space by a depth of 2000 m (c=1.0). As
summarized in Table S2 and Figure S1, variations in the value of c result in a range of estimates
of compaction tilt (Table S2) and impact the estimated monoclinal relief increase (Figure S1) in
a non-linear fashion, because of the combined effects of the exponential porosity-depth function
and the spatial variations in thicknesses of L and M and the Cenozoic units which they overlie.
For these tests (“c” varies while sonic velocity is held constant, for the Toconao-1 velocity
model), the standard deviations on estimated monoclinal relief growth of individual horizons in
the three seismic lines range from 137 m to 384 m, constituting uncertainties between about 4%
and 15%. Based on Flint [1987] and Nadon and Issler [1997], we anticipate that a reasonable
choice may be c~0.4, which creates a decrease in inter-granular space to values of 16% and 7%
at depths of 2000 m and 4000 m, respectively. In the absence of any in situ constraints on intergranular cement-plus-void volume for the subsurface of the eastern Salar de Atacama we do not
argue that this is the correct compaction relationship, but use c=0.4 as a basis for comparison of
the set of velocity models.
S3. Calama Basin
S3.1. Stratigraphy
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Middle and Upper Miocene strata of the eastern part of the Calama basin are alluvial fan
deposits [May et al., 1999; Blanco, 2008]. Unlike in the Pampa del Tamarugal, in the externallydrained Calama basin there are no modern alluvial fans fed from the same catchment basins as
were the Middle and Upper Miocene fan deposits. Given the combination of climate change
[Rech et al., 2006, 2010] and lack of modern analogue fans, we cannot estimate the depositional
inclinations of the Miocene alluvial deposits with sufficient confidence to proceed with an
estimation of post-depositional rotation of these strata.
S3.2. Short Wavelength Structures and Non-Tectonic Rotation
There has been less deformation of the Calama basin during the Middle Miocene to
Holocene than in either the Pampa del Tamarugal or the Salar de Atacama basins. North-trending
strike-slip deformation near the western limit of the Calama valley is extensively documented
because of its relationship to copper mineralization, but associated vertical displacements
affected local topography rather than basin wide tilt [Tomlinson and Blanco, 1997; Tomlinson et
al., 2001]. A small number of tectonic faults and folds of local importance within the center of
the Calama basin have been mapped by various workers [e.g., Tomlinson and Blanco, 1997;
Jordan et al., 2006; Houston et al., 2008; Blanco and Tomlinson, 2009] and illustrated by
Ramirez [1979] and Houston [2004, 2007] in the eastern part of the basin. Among this set of
faults, the early Neogene faults near the center of the basin trend NNE to NE and accommodated
hundreds of meters or more of down-to-the-east normal offset [Jordan et al., 2006]. Near the
eastern basin margin, NW-trending normal faults displace a ~9 Ma unit by less than 30 m
[Houston, 2007]. In sum, the known faults and folds either lie west of the region that we analyze
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for rotation of the eastern branch of the Western Andean slope, or offset units in a direction
opposite to that of the long-wavelength uplift, or are of small magnitude.
Of direct importance as a potential source of westward tilt of the eastern part of the
Calama basin is a NNE-striking, east-vergent reverse fault system located along the east margin
of the Tuina-Aiquina basement high (Fig. 6c, buried dashed fault, modified from Ramirez
[1979]). The N25E-trending Tuina-Aiquina highland separates the Calama basin from the Salar
de Atacama basin and from the steepest part of the eastern branch of the Western Andean Slope
(Fig 6c). The Tuina-Aiquina highland has acted as a hydrographic divide between the Calama
and Salar de Atacama basins since at least the Oligocene [e.g., May et al., 1999; Blanco, 2008]
and the fault system along its east margin has been active during various episodes, including an
inferred displacement of 100’s of meters during the Late Miocene [Blanco, 2008]. There is very
little information on the geometry, style or timing of this important fault set and associated folds,
which are extensively buried by young volcanic rocks. Exposures of pre-Pliocene strata suggest
that the trace of this blind fault system is ~10 km east of the eastern extent of Calama basin strata
(Fig. 6c), and roughly at the westernmost position possible for the upper hinge of the eastern
limb of the Western Andean Slope. Vertical displacement on this fault might have contributed to
post-20 Ma relief development of the NNE-trending flank of the Calama basin.
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References cited in Supplementary Data that are not already given in the primary reference list:
Bouzari, F., Clark, A.H., 2002. Anatomy, evolution and metallogenic significance of the
supergene orebody of the Cerro Colorado porphyry copper deposit, I region, northern Chile.
Economic Geology 97, 1701 – 1740.
Damanti, J. F. and T. E. Jordan (1989) Cementation and compaction history of synorogenic
foreland basin sedimentary rocks from Huaco, Argentina, AAPG Bull., 73, 858-873.
de Silva, S. L. (1989b) Geochronology and stratigraphy of the ignimbrites from the 21°30′S to
23°30′S portion of the Central Andes of northern Chile, Journal of Volcanology and
Geothermal Research, 37, 93-131.
Flint, S. (1987) Diagenesis of Tertiary playa sandstones of Northern Chile; implications for
Andean uplift and metallogeny, Sedimentology, 34, 11-29.
Houston, J. (2004) High-resolution sequence stratigraphy as a tool in hydrogeological
exploration in the Atacama Desert, Q. J. Eng. Geol. Hydrogeol., 37, 7-17.
Houston, J. (2007) Recharge to groundwater in the Turi Basin, northern Chile: An evaluation
based on tritium and chloride mass balance techniques, J. Hydrol., 334, 534-544.
Lindsay, J. M., S. de Silva, R. Trumbull, R. Emmermann, and K. Wemmer (2001), A
reevaluation of the stratigraphy and volcanology of one of the world’s largest resurgent
calderas, Journal of Volcanology and Geothermal Research, 106, 145-173.
Mpodozis, C., N. Blanco, T. E. Jordan, and M. C. Gardeweg (2000) Estratigrafía, eventos
tectónicos y deformación del Cenozoico tardío en la región norte de la cuenca del Salar de
Atacama: la zona de Vilama-Pampa Vizcachitas, Congreso Geológico Chileno, 9, 598-603.
Muñoz, N. and P. Sepulveda (1992) Estructuras compresivas con vergencia al oeste en el borde
oriental de la Depresion Central, norte de Chile (19 degrees 15'S), Revista Geologica de
Chile, 19, 241-247.
Nadon, G. C. and D. R. Issler (1997) The compaction of floodplain sediments: timing, magnitude
and implications, Geoscience Canada, 24, 37-44.
Ramírez, C. (1979) Geología del Cuadrángulo Río Grande y sector nororiental del Cuadrángulo
Barros Arana, Provincia El Loa, II region, Geologist thesis, 139 pp., Universidad de Chile,
Santiago, Chile.
Vergara, M., C. Marangunic, H. Bellon, and R. Brousse (1986) Edades K-Ar de las ignimbritas
de las Quebradas Juan de Morales y Sagasca, norte de Chile, Serie Comunicaciones,
Departamento de Geología, Facultad de Ciencias Físicas y Matemática, Universidad de
Chile, 36, 1-7.
15
Figure 1: Uncertainty on monoclinal uplift produced by uncertainty of burial compaction. The
apparent structural relief formed across the monocline east of Salar de Atacama is shown for a
suite of plausible burial compaction histories. For three seismic lines (1g006, 1g002, and 1g16b),
and two horizons on each of those lines (base of M, base of L), differential compaction of the
east and west ends of the inclined segments were calculated using Sclater and Christie’s (1980)
exponential compaction function, fz = foexp(-cz), for a broad range of values of “c.” Z is depth
below the surface, fz is the porosity at burial depth z, fo is the porosity that the sediments
possessed at the surface prior to burial, and “c” is a term that describes the shape of the curve of
decrease in porosity with depth. A “c” value of 0 corresponds to no compaction, and Sclater and
Christie (1980) document “c” values between 0.27 and 0.51 for marine siliciclastic strata of
various textures. The analysis shows most broadly that significant burial compaction would
cause a small degree of rotation in the same direction as the tectonic monocline, and correction
of the end-product inclination for that compaction diminishes the perceived tectonic monoclinal
uplift. In more detail, we find the maximum compaction affect for a “c” value between 0.3 and
0.5 for the suite of seismic lines and horizons, hence we use a “c” value of 0.4 in corrections for
both the Salar de Atacama and Pampa del Tamarugal data and in other tests of uncertainty. If the
uncertainty is considered to be the percentage of difference between the calculated structural
relief if there was no compaction compared to the calculated structural relief if “c” = 0.4, divided
by the structural calculated relief if there were no compaction, the mean maximum uncertainty
on tectonic monoclinal uplift due to lack of information about compaction is 19%.
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