Lithospheric thinning and dynamic uplift effects during slab window formation, southern Patagonia (45˚-55˚ S)

Journal Pre-proof Lithospheric thinning and dynamic uplift effects during slab window ˚ formation, southern Patagonia (45˚-55S) ´ Pilar Avila (Conceptualization) (Methodology) (Software) (Validation) (Formal analysis) (Investigation) (Resources) (Writing original draft) (Writing - review and editing) (Visualization) ´ (Supervision) (Project administration), Federico M. Davila (Conceptualization) (Methodology) (Software) (Validation) (Formal analysis) (Investigation) (Resources) (Writing - original draft) (Writing - review and editing) (Visualization) (Supervision) (Project administration) (Funding acquisition)

PII:

S0264-3707(19)30149-8

DOI:

https://doi.org/10.1016/j.jog.2019.101689

Reference:

GEOD 101689

To appear in:

Journal of Geodynamics

Received Date:

6 June 2019

Revised Date:

15 November 2019

Accepted Date:

3 December 2019

´ ´ Please cite this article as: Avila P, Davila FM, Lithospheric thinning and dynamic uplift effects ˚ Journal of Geodynamics (2019), during slab window formation, southern Patagonia (45-˚55S), doi: https://doi.org/10.1016/j.jog.2019.101689

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Lithospheric thinning and dynamic uplift effects during slab window formation, southern Patagonia (45˚-55˚ S)

Pilar Ávila and Federico M. Dávila

CICTERRA, CONICET - Facultad de Ciencias Exactas, Físicas y Naturales, Universidad Nacional de Córdoba, Córdoba 5016, Argentina;

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Corresponding author: [email protected]

ABSTRACT

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The southernmost South America has been affected by the subduction of an oceanic

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seismic ridge, which began to subduct below southern Patagonia at ~14 Ma. This scenario led to the formation of a slab window, which is still active and where hot buoyant asthenospheric

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mantle produced thermal anomalies and modifications in the lithospheric thicknesses. Meanwhile, from the Patagonian Andes to the Atlantic coast, an outstanding regional surface

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uplift took place, conducting to a moderated-elevation plateau formation. In this work we analyzed the causes of this long-wavelength surface elevation change using residual

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topography and uplift rate calculations, considering paleo-lithospheric states. We achieved this considering that the study area underwent a lithospheric thickness changes through time, before and after the slab window formation. This allowed us to estimate the isostatic and dynamic adjustment over time and their influences on the surface elevation changes. These rates were compared with geological and stratigraphic observations derived from a key elevation marker bed: The modern altitudes of Oligo-Miocene marine strata top, originally deposited close or below sea level, and placed at Present at hundreds of meters above sea level. Our residual topography calculations, that result from comparing isostatic and observed topographic,

indicates the dynamic topography contribution in the study region was very minor (if so) to null. The isostatic uplift, in turn, shows a remarkable fitting with the reconstruction of marine marker bed, suggesting a causative relationship. We can assert that Patagonian lithospheric thinning, particularly of the lithospheric mantle, by slab window formation, would have been enough to reproduce the modern elevations and surface uplift across the plateau from the Miocene to Present day. Our work opens a question on the model that connect slab windows, asthenospheric upwelling flows with surface uplifting. However, as shown by recent works,

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lithospheric changes might trigger small convection cells, which might produce local and small

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dynamic topography.

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Keywords: slab window; dynamic topography; lithosphere; residual topography;

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isostasy; southern Patagonia

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1. Introduction

The Earth’s surface topography is governed by different processes that act at different scales (Artemieva, 2011). The most influential is isostasy, which assumes that the different

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lithospheric and asthenospheric forces are in gravitational equilibrium (Lachenbruch and Morgan, 1990). Any crustal or/and lithospheric mantle change as well as surface erosion,

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sedimentation and/or rock strength modification might therefore trigger isostatic adjustments and, consequently, drives surface uplift or subsidence. The asthenospheric mantle can also play an important role in surface topography by deforming the lithosphere at regional and continental scales (Lithgow-Bertelloni and Richards, 1998 among others), altering the isostatic equilibrium. This long-wavelength deformation process is caused by the vertical stresses arising from viscous flow in the mantle in response to its 3D density structure, known as dynamic topography (see Hager et al., 1981). Nevertheless, this contribution is difficult to

detect and measure, given that total topography is (as stated above) overprinted by isostatic topography. The computation of the isostatic surface topography requires of knowing mainly crustal and lithospheric mantle thicknesses and densities. Different databases have made a large effort to compile this information at global, regional and local scales, which are based on diverse and contrasting geophysical surveys and methodologies that have different uncertainties and precisions (Artemieva and Mooney, 2001; Simmons et al., 2010; Steinberger and Becker,

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2016). On the other hand, dynamic topography are based on two different and contrasting

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approaches: (i) forward and adjoint models i.e., computing mantle convection (e.g., LithgowBertelloni and Richards, 1998; Liu et al., 2008; Flament et al 2013; among many others) and/or

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(ii) residual topography, i.e., removing the isostatic component to the total topography (e.g., Faccenna et al., 2014; Artemieva and Vinnik, 2016; Artemieva 2007). Even though mantle

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flow numerical forward models have substantially improved (see Liu and Zhong, 2016), they

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are still based on a large amount of assumptions. This would have conducted many authors to test mantle dynamic topography by using residual topography (e.g., Dávila and LithgowBertelloni, 2013; e.g., Faccenna et al., 2014; Flament et al., 2013; among others).

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Southernmost South America, particularly southern Patagonia (with an area of ~500,000 km2), is an attractive geodynamic setting to study topography evolution during the Miocene to

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Present day. This region has been affected by the subduction of a seismic ridge (the South Chile spreading Ridge), which began to subduct below southern Patagonia at ~14 Ma. This led to the formation of an asthenospheric or slab windows (Thorkelson, 1996), where hot buoyant asthenospheric mantle produced thermal anomalies and consequent changes in the LAB (Lithosphere Asthenosphere Boundary) depth (e.g., Ávila and Dávila, 2018). Along the Patagonian Andes and foreland, an abrupt and major increase of both summit elevations and local relief along the orogenic crest occurred immediately inland of the Chile Triple Junction

(CTJ), where the Chile Ridge collides with the South American continent. Particularly, toward the distal foreland and Atlantic coast, a moderated-elevation (1000-500 m asl) plateau developed. Recent studies have proposed a long-wavelength pulse of dynamic uplift, ensuing exhumation above the Patagonian slab window given that these gap might drive mantle upwelling (Dávila et al., 2018; Guillaume et al., 2013, 2010, 2009). These predictions come mainly from analog and forward numerical modeling (e.g., Guillaume et al., 2013, 2010). However, other works (e.g. Georgieva et al., 2016; Thomson et al., 2010) have questioned the

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dynamic topographic connection based mainly on low-temperature thermochronology. While

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Thomson et al. (2010) suggested a regional glacial protection from erosion, leading to constructive growth in orogen height and width in the southern Patagonian Andes, Georgieva

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et al. (2016) defined a pulse of glacial incision coeval with neotectonic activity since 3-4 Ma, supported by structural and geomorphological evidence.

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In this work we analyzed the causes of long-wavelength (plateau) uplift in southern

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Patagonia using residual topography and uplift rate calculations. We considered not only the current geodynamic setting but also the lithospheric changes through time from middle Miocene to Present day, i.e., before and after the slab windows formation (Breitsprecher and

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Thorkelson, 2009). This allowed us to estimate the isostatic and dynamic adjustment over time and their influences on surface uplift or subsidence. These rates are compared with geological

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and stratigraphic observations derived from a key elevation marker bed. We use the modern altitudes of Oligo-Miocene marine strata top (originally formed close or below sea level, and today placed hundreds of meters above sea level). This will assist us quantifying the major controls on topography during ocean ridge subduction and slab window formation by alteration of the lithospheric structure (e.g., Ávila and Dávila, 2018) and/or asthenospheric mantle convection (Dávila and Lithgow-Bertelloni, 2013; Guillaume et al., 2009).

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2. Geological and tectonic setting

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Figure 1 Map of southern Patagonia showing the location of the main stratigraphic, tectonic and volcanic features, referenced in our work. Austral and Golfo de San Jorge basins are in light grey. Cenozoic basalts of the Patagonian plateau lavas are shown in dark grey (Boutonnet et al., 2010; Espinoza et al., 2010; Gorring et al., 2003); seismic stations location are shown as diamonds (Robertson Maurice et al., 2003). The Andean thrust fronts (easternmost thrusts) are represented by the black solid lines. The Oligocene-early Miocene marine beds are shown in orange. CTJ: Chile Triple Junction. 1 and 2 are the northern and southern margins of the slab window, (taken form Breitsprecher and Thorkelson, 2009). FTS: Fagnano transform system.

The study region is located in southern Patagonia, southernmost South America (southern

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Argentina and Chile), between 45° and 55° S, embracing from the southernmost Andean belt, to the west, to the foreland and passive Atlantic platform, to the east (Fig. 1). Southern Patagonia is tectonically bounded to the west by the Nazca and Antarctic oceanic plates, which are in turn separated by the oceanic Chilean Ridge. This plate configuration generated a triple junction against the continent at ~46.5 ° S (Fig. 1). To the north of the junction the subduction of Nazca plate is “normal”, in terms of dip angle and extent within mantle, whereas to the south the Antarctic plate subduction is only 100 km beneath the South American plate (Breitsprecher and Thorkelson, 2009). The oceanic Chilean Ridge as well as Chile Triple Junction (CTJ)

would have stayed at its present location for the last ~4 my (Breitsprecher and Thorkelson, 2009). Both tectonic features migrated ~1000 km northward, from 54° S, since 14 Ma (Middle Miocene), as ridge segments subparallel to the trench collided with the subduction margin (Breitsprecher and Thorkelson, 2009). This evolution led to the formation of an asthenospheric, where hot buoyant asthenospheric mantle produced thermal anomalies and consequent changes in the LAB depth (Fig 2.a Ávila and Dávila, 2018). Spatial voids in a subduction were commonly associated with upwelling flows and continental dynamic uplift, not only in

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Patagonia (see Dávila et al., 2018; Guillaume et al., 2013, 2009) but also in other regions

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(western United States, see Becker et al., 2014). The definition of the Patagonian slab window edges and area are shown in Figure 1 (based on petrological studies, basalt geochemistry and

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paleogeographic reconstructions, see compilation in Breitsprecher and Thorkelson, 2009). Asthenospheric window formation since the middle-late Miocene in Patagonia was

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contemporaneous with: (1) a major exhumation event and beginning of glaciations in Patagonia

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along the Cordillera (Georgieva et al., 2016; Guillaume et al., 2013; Thomson et al., 2010), (2) plateau uplift across the foreland (Dávila and Lithgow-Bertelloni, 2013), and (3) absence or poor crustal thickening (Lagabrielle et al., 2004). The generalized uplift of the foreland (no

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crustal deformation) is evidenced by different geological observations. In the Oligocene-early Miocene, immediately previous to plateau uplift, Patagonia was under sea level as suggested

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by the large distribution of marine beds (Centinela, Monte Leon, Rocas Verdes and 25 de Mayo Fms, also known as “Patagoniano, Cuitiño et al., 2016; Encinas et al., 2018) (Fig. 1). These strata were stratigraphically followed by a regional erosive discordance and deposition of lower-middle Miocene alluvial to estuarine deposits (Cuitiño et al., 2016; Encinas et al., 2018). The Santa Cruz Fm (and correlative units) have been related to the main phase of Cenozoic (Andean) deformation and surface uplift along the Patagonian Cordillera (see Fig. 1). This thick continental sedimentary pile (between 500 and 1000 meters) was constrained between ~19 Ma

(youngest marine Patagoniense bed, Cuitiño et al., 2016) and ~12 Ma (oldest Miocene basalt resting on this unit; Folguera et al., 2018). The Santa Cruz Fm is covered by relatively thin (<100s meters thick) Plio-Pleistocene tillites (Rabassa et al., 2005 and references therein) and the Rodados Patagonicos, interpreted as foreland bypass gravel sheets (Ghiglione et al., 2016), likely associated with the main plateau foreland uplift of Patagonia (see Guillaume et al., 2009). These gravelly beds are interlayered by OIB basalts, considered an independent evidence source of the slab windows formation since the middle-late Miocene (e.g. Boutonnet et al.,

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2010; Espinoza et al., 2010; Gorring et al., 2003). These upper Miocene to Present basalts are,

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in turn, subhorizontal and cover the easternmost Andean thrust front and most of the foreland areas (see Fig. 1), supporting no major crustal deformation events. Furthermore, in the

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easternmost part or pericratonic Patagonian foreland, an areally extensive and transient uplift has been documented by elevated Pleistocene marine terraces along the Atlantic coast. They

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can locally reach 180 m asl (Pedoja et al., 2011). All these evidences (non-tectonic uplift of the

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Patagonian foreland and its correlation with the arrival of the oceanic Chile Ridge) drove different authors to propose that the main plateau uplift, locally near ~1 km, in southern Patagonia could have been conducted by slab-window driven upwelling mantle flows (dynamic

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uplift) (Dávila et al., 2018; Guillaume et al., 2009; Flament et al., 2015). In this contribution we calculated residual topography and uplift rates to analyze the influence of deep processes

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(dynamic topography and changes in the lithospheric mass, e.g., lithospheric thinning by heat flow increasing, Ávila and Dávila, 2018) on surface uplift across the foreland during ocean ridge subduction and slab window formation.

3. Residual topography 3.1 Method and model

The Earth’s surface topography is mainly governed by isostatic (𝑇𝑖𝑠𝑜 ) and dynamic (𝑇𝑑 ) forces (Fig 3a). The isostatic elevations depend on the mass balance between the lithosphere and asthenosphere. Thus, the crustal and lithospheric mantle thicknesses, as well as densities within the lithosphere and asthenosphere, are relevant values to take into account in isostatic calculations. Dynamic topography in turn, results from viscous forces driven by mass anomalies generated within the mantle (see Lithgow-Bertelloni and Richards, 1998). On the base of these two main forces, surface topography (or total topography, 𝑇𝑡𝑜𝑡𝑎𝑙 ) may be written

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as follow:

(1)

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𝑇𝑡𝑜𝑡𝑎𝑙 = 𝑇𝑖𝑠𝑜 + 𝑇𝑑

The mean isostatic elevation (𝑇𝑖𝑠𝑜 ) is argued to be mainly dominated by distribution of

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crustal masses (roots and anti-roots). Zoback and Mooney (2003), however, have demonstrated

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the important influence of the lithospheric mantle in isostatic calculations. Then, 𝑇𝑖𝑠𝑜 can be estimated as follow:

𝜌𝑎

𝜌𝑚𝑙 −𝜌𝑎

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𝜌𝑐 −𝜌𝑎

𝐻𝑐 +

𝜌𝑎

𝐻𝑚𝑙 − 𝐻0 )

(2)

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𝑇𝑖𝑠𝑜 = 𝛼(

where ?? is 1.45 for calculation under sea level or 1 for continent: 𝜌𝑐 , 𝜌𝑚𝑙 and 𝜌𝑎 are

densities of the crust, mantle lithosphere and asthenosphere, respectively; 𝐻𝑐 and 𝐻𝑚𝑙 are crustal al lithospheric mantle thicknesses, respectively; 𝐻0 is the buoyant height of sea level relative to the hypothetical free surface of the asthenosphere that is, the “asthenosphere geoid” (Fig 3b). Neglecting the flexural effects due to lithosphere elasticity (which is reasonable at wavelengths >200 km), the difference between observed (or total) and isostatic topography

might be considered as a good proxy and indicative of the dynamic topography (see Flament et al., 2013). We calculate our residual topography considering that:

𝑇𝑟 = (𝑇𝑑 ) = 𝑇𝑡𝑜𝑡𝑎𝑙 − 𝑇𝑖𝑠𝑜

(3)

Positive 𝑇𝑟 values evidence total topographies larger than isostatics, suggesting the presence of sublithosperic downwarping forces to match the supposed excess of topography.

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Negative values would indicate the opposite, whereas zero values (or close to zero) are

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evidence of equilibrium, when no additional forces are required to account for the observed

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topography.

3.2. Input parameters

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In comparison with other regions of South America, Patagonia has a rather poor

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geophysical database. For our calculation we used two different types of datasets: 1) local and regional data (see below) and 2) a global model for the Earth's lithosphere and upper mantle LithosRef18 (Afonso et al., 2019). LithoRef18 is a global model for the Earth's crust and

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lithospheric mantle obtained through a formal joint inversion of 3D gravity anomalies, geoid height, satellite-derived gravity gradients and absolute elevation complemented with seismic,

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thermal and petrological information. The model includes crustal thickness, average crustal density, lithospheric thickness, and depth-dependent density of the lithospheric mantle, with a surface discretization of 2ºx2º.

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Figure 2 Input parameters used in residual topography computations. a) Surface heat flow distribution (taken from Ávila and Dávila, 2018). White dots show the location of the analyzed boreholes. Black dots shows the position of boreholes (Rodriguez and Littke, 2001; Sachse et al., 2016), where we estimated paleo-heat flows. b) Thermal LAB thicknesses (cf. Ávila and Dávila, 2018). c) Differences between our LAB thermal model and LAB LithosRef18. d) Crustal thicknesses from van der Meijde et al. (2013). e) Map showing the differences between (d) and crustal thicknesses from LithosRef18 (red-white scale). The lines show the difference between (d) and local values from Alcacer et al. (2017). f) In situ crustal and g) lithospheric mantle densities from LithosRef18.

Crustal thickness values (𝐻𝑐 ) were taken from GMSA12 regional crustal model (van der Meijde et al., 2013, see Fig 2c). This model derives crustal thickness from satellite gravity data.

It utilizes the combined gravity model EIGEN-6C, which is composed of GOCE. The GMSA12 model has been compared with point constraints on crustal thickness for South America and the oceanic part surrounding the continent. A total of 736 comparisons have been made (however, none in Patagonia). The point constraints have, on average, an error of ±3 km (Assumpçăo et al., 2013) and the uncertainty related to model parameters is also around ±3 km. The values derived from this database correlates rather well with LithoRef18 model (with differences up to 6 km, see Fig 2e), available regional (Chulick et al., 2013; Feng et al., 2007)

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and local data (Alcacer et al., 2017; Robertson Maurice et al., 2003) from Patagonia. It is

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important to notice that the difference between these regional and local datasets (Alcacer et al., 2017; Robertson Maurice et al., 2003) is ± 10 km in southern Patagonia (Fig 2c). We used this

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difference as the uncertainty for the crustal thickness in the study region.

As mentioned previously, our modeling is based on the assumption that the isostatic

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balance is achieved at the LAB, and thus knowledge on the lithosphere thickness is crucial. We

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used Ávila and Dávila (2018) LAB model, which is the first regional-scale distribution of the LAB in southern Patagonia. In this model the thermal structure of the lithosphere is based on the analysis of regional surface heat flow data measured in a large number of 1000 m to 6000

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m deep boreholes from two of the most productive petroleum basins in Argentina: the Golfo de San Jorge and Magallanes-Austral basins (Fig. 1 and Fig 2a). In this model, the base of the

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lithosphere is defined as the base of the thermal boundary layer, represented locally by the intersection of conductive geotherm and mantle adiabat with temperature of 1300 °C (the detail related to model assumptions and parameterization can be found in Ávila and Dávila, 2018). These results evidence an eastward increasing of the lithospheric thicknesses between 25 and 55 km in southernmost Patagonia (Fig 2d). This dataset is in agreement with local studies (four local stations located in the southwesternmost part of the window, see Fig. 1) of LAB depth inferred from regional waveform inversion (Roberson Maurice et al., 2003). Roberson Maurice

et al. (2003) inverted regional seismograms recorded by a Seismic Experiment in Patagonia and Antarctica. The inversion included anisotropy by solving for separate SV and SH structures in the upper mantle. Results show the strongest anisotropy is localized in a lithospheric lid shallower than 65 km depth, overlying a pronounced low-velocity zone. The lack of thermal equilibrium in the northern region does not allow the thermal LAB to be calculated using the steady-state equation. Nonetheless, the authors propose lithospheric thicknesses similar to those reported to the south, in the Austral basin, given that the Golfo de San Jorge area would

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have been embraced by the asthenospheric window since 4 Ma (Breitsprecher and Thorkelson,

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2009). For this reason we fix LAB depth for the northern region with the same values calculated for the south. This dataset in interpolated with thermal LAB data calculated by Tassara and

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Echaurren (2012). Our LAB input has similar values to those proposed in LithoRef18, with differences that do not exceed 15 km (Fig 2f). We used this value as uncertainty for thermal

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LAB since they have similar values proposed by other thermal model (ca. 25% for the

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lithospheric thickness, Artemieva, 2003; Artemieva, 2006).

The density structure of the entire continental lithosphere is poorly known due to a large uncertainty in the genesis and composition. In this contribution in situ density values for the

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crust and lithospheric mantle were taken from LithosRef18 (Afonso et al., 2019). In these model, density is assumed to be a function of pressure and temperature. A comparison with

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higher-resolution density models (for the crust and lithospheric mantle) in different parts of the world (e.g. Hasterok and Gard 2016; Alghamdi et al. 2018; Yegorova et al. 2013; among others) shows a close agreement with densities values (crustal and lithospheric mantle) of LithosRef18 with most of the differences being within 0.05 g/cm3. Given that no local values exist in southern Patagonia, both densities are assumed to have an error of 0.1 g/cm3. Density of asthenosphere was considered as constant = 3.235 g/m3 at in situ mantle conditions (T =

1300°C), which corresponds to density of 3.39 g/m3 at SPT (Standard conditions for temperature and pressure) conditions (Cherepanova and Artemieva, 2015). 𝐻0 can be obtained using a mid-ocean ridge as a reference density column (Lachenbruch and Morgan, 1990). This estimation assumes that the oceanic crust rests directly on asthenosphere along the oceanic ridge crests. On the base of this idea, Lachenbruch and Morgan (1990) proposed an average value of 2.4 km assuming the Moho depth at 5.5 km, the water depth at 3.5 km, and the average crustal density of 2.8 g/m3. Although the chosen values

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correspond to typical observations over mid-ocean ridges, they do not guarantee that regions

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with these parameters of crustal structure and bathymetry are in isostatic equilibrium as required for our calculation. More recently, Cherepanova and Artemieva (2015) re-evaluated

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the constant 𝐻0 by choosing those segments of mid-ocean ridges isostatically compensated (for example the triple junction of the Pacific, Nazca and Cocos plates). Given the uncertainty in

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the Moho depth and the average crustal density, they calculated a family of possible solutions

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for the constant by varying the crustal density, Moho depth, and bathymetry. These author proposed a 𝐻0 = 4.25 km that corresponds to a oceanic crust 6 km thick, density 2.75 g/m3 and

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bathymetry 3 km.

3.3. Sensitivity tests

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We run sensitivity tests to examine how uncertainties in the model parameters propagate

to uncertainty in the final residual topography calculation. The isostatic contribution of the crust and lithospheric mantle is calculated directly from crustal and lithospheric mantle thickness and density data (Eq. 3). The results are summarized in Table 1. The tests are calculated and shown for the maximum values of the estimated uncertainties for each model parameter. On the whole, in the most unfortunate case, when all uncertainties sum up, the total residual topography uncertainty is ca. 0.9 km. Given that it is unlikely that, at any location, the

errors in all model parameters work in the same direction, we estimated the uncertainty of the final residual topography model as better than 0.5 km.

Reference model Crustal Thickness

LAB thickness Crustal density (in situ) Lithospheric Mantle density (in situ) Asthenosphere density (in situ)

33 43 2.820

Parameter variation with respect to reference model +10 km +15 km +0.1 g/m3

Residual Topography (km) 0.7 less 0.06 less 0.2 less

3.270 3.235

+0.1 g/m3 +0.05 g/m3

0.5 less 0.03 less

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Table 1 Sensitivity test. Reference model is the average value of each parameter calculated for the entire study region.

of ro -p re lP na Jo ur Figure 3. a) Three mass columns (in grey) illustrating from left to right: Isostatic topography (Tiso) with no dynamic topography component; dynamic upwelling (red arrow) and downwelling (blue arrow). Ttot: total topography; Td: dynamic topography; SL: sea level. b) Lithosphere buoyancy model after Lachenbruch and

Morgan (1990). Mass columns are assumed to “float” on an asthenosphere of density ρa. The first two columns, from left to right, show a general elevation (Tiso) and lithospheric thickness (L), compound of crust and mantle with densities ρc and ρml, respectively. The third column, to the right, is a mid-ocean ridge, where ocean crust is assumed to rest directly on asthenosphere. This column is taken as a reference column, used to calculate the height of the asthenosphere geoid (H0). ρw: density of seawater (modify from Zoback and Mooney et al., 2003) . c) Surface subsidence (𝑉𝑖1 ) and uplift (𝑉𝑖2 ) generated by a crustal thickness (Hc) variation through time (notice that the lithospheric mantle thickness (Hml) remains constant over time and the lithosphere is under isostatic compensation). t: time. d) Surface subsidence (𝑉𝑖1 ) and uplift (𝑉𝑖2 ) generated by a lithospheric mantle thickness variation through time (notice that the crustal thickness remains constant over time and the lithosphere is under isostatic compensation). e) Surface uplift (𝑉𝑑1 , 𝑉𝑑2 and 𝑉𝑑6 ) or subsidence (𝑉𝑑3 , 𝑉𝑑4 and 𝑉𝑑5 ) generated by changes of dynamic topography over time.

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4. Uplift rate 4.1 Method and model

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The vertical displacement of Earth surface through time (𝑉𝑠 see Eq. 4) is mainly driven by isostatic (𝑉𝑖 , see Fig 3c-d) and dynamic topography changes (𝑉𝑑 , see Fig 3e) through time

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(∆𝑡). Therefore, the uplift (or subsidence) rate (𝑉𝑠 ) can be written as the derivative of total

∆𝑡

= 𝑉𝑠 = 𝑉𝑖 + 𝑉𝑑

𝑉𝑖 = (

𝜌𝑎 −𝜌𝑐 ∆𝐻𝑐

)

+(

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∆𝑇𝑡𝑜𝑡

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topography (𝑇𝑡𝑜𝑡𝑎𝑙 ) in time (see Eq. 3).

𝜌𝑎

∆𝑡

𝜌𝑎 −𝜌𝑚𝑙 ∆𝐻𝑚𝑙 𝜌𝑎

)

∆𝑡

𝑎𝑛𝑑 𝑉𝑑 =

∆𝑇𝑑 ∆𝑡

(4)

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These dynamic and lithospheric changes can be driven by a complex interaction between, crustal tectonics, thermal state of the lithosphere among others. In consequence, to calculate

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the amount and direction of Earth surface displacement, at least two successive topographic states must be considered (Fig 3c-d-e). Therefore, if we aim to test the influence of dynamic topography and lithospheric thinning in surface uplift in Patagonia since the arrival of the asthenospheric window, two scenarios must be analyzed: before and after the slab windows formation. We calculated 𝑉𝑠 and 𝑉𝑖 (like in section 3.1 for estimating residual topography), 𝑉𝑑 results from subtracting the isostatic component (𝑉𝑖 ) from uplift (or subsidence) rates (𝑉𝑠 ) (see above).

4.2. Oligo-Miocene topography and estimation of 𝑉𝑠 . Considering that the southern Patagonian foreland was at sea level in the late Oligocene - early Miocene (OM from now), and at Present-day it places hundreds of meters above, we used this marine record as an elevation marker to estimate surface uplift since the OM to Today. It is important to mention that sea level changes were <100 m since the OM (cf. Haq et al., 1987; Miller et al., 2005) i.e., much lower than the Present-day elevations of the OM marine

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units. Moreover, no crustal deformation has been reported in the Patagonian foreland, eastward

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from the Andean thrust front, since the OM (Lagabrielle et al., 2004 see Fig 1). Therefore, any elevation change might be connected with subcrustal forces. Given that the OM levels have not

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been used for topographic analysis, we first had to map the altitude top (or uppermost section) of these units (known with different names, Centinela, Monte Leon, Rocas Verdes and 25 de

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Mayo Fms, or “Patagoniano”, see Cuitiño et al., 2016; Encinas et al., 2018 among others Fig

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1) in the study region. For this reconstruction (Fig 1 and Fig 5) we used five databases: (a) Panza (2003) synthesis, (b) Argentine Geological Survey geological maps (Ardolino et al., 2003; Dal Molin et al., 1998; Escosteguy et al., 2003; Giacosa et al., 1998; Panza et al., 2003;

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Sciutto et al., 2000), (c) Landsat images (taken from National Oceanic Atmospheric Administration) to reference the units with better accuracy and (d) 90 m SRTM digital elevation

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model, from where we extracted the altitude values for the surface reconstruction. In those regions where the OM markers place in subsurface, we used information from (e) oil well data reports (see Ávila and Dávila, 2018). When possible, we mapped the boundary between the OM marine beds and Santa Cruz (and correlative) Fm (i.e., OM top). In those regions where the OM top is not present, we selected the highest sections of the OM beds. This approach would not have large errors across the foreland, because the marine beds are relatively thin (100-50 m thick). Toward the Cordillera, where the OM can be >500 m thick (see Encinas et

al., 2018), we only picked up the OM tops to avoid large errors. After mapping different elevation points, we reconstructed the modern surface elevation of the OM boundary by making a planar tendency from outcrops and subsurface data. This selection agrees with an originally flat accumulation surface (siliciclastic shallow-marine platform environments), was tilted during a subsequent non-tectonic uplifting event. We analyzed data inside and outside the

4.2 Paleo-lithosphere calculation and estimation of 𝑉𝑖 .

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prediction bounds and the goodness-of-fit for the planar tendency.

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The estimation of 𝑉𝑖 requires of knowing two successive lithospheric states. As stated above, we compared the Modern and Miocene scenarios (i.e., before and after slab windows

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formations). While the Modern isostatic state was developed above (see also Ávila and Dávila, 2018), the Miocene isostatic computations require to estimate a paleo-lithospheric thickness.

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For this purpose, we used a paleo-thermal approach. Variations on Earth surface heat flow can

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be interpreted in terms of the thermal structure of the lithosphere (Artemieva, 2011). In stable areas of continents, where no tectonic activity has been experienced for several million years, the thermal structure can approximated by the steady state solution of the thermal conductivity

𝐴(𝑧)

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𝑑2 𝑇

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equation (see Artemieva and Mooney, 2001) The thermal conductivity equation is given by:

𝑑𝑧 2

= 𝑘(𝑇)

(5)

with boundary conditions at surface:

𝑇|𝑧=0 = 𝑇0

𝑑𝑇

𝑄0 = −𝑘 𝑑𝑧

(6)

(7)

where 𝑄0 is the near surface heat flow, 𝑇0 is surface temperature, T is temperature a function of depth, k is thermal conductivity, and 𝐴 = 𝐴(𝑧) is the heat production as a function of depth. The solution of Equation 5 allows determining the distribution of temperature with depths. The 1300°C isotherm corresponds to the LAB depth. These calculations require to define the surface heat flow and the distribution of thermal parameters (thermal conductivity

Surface temperature

0 ºC

Q0

Surface heat flow

50-53 mW/m2

Radioactive heat production equation Surface radiogenic heat production rate per unit mass

upper crust

𝐴0

hr

Conductivity of the upper continental crust equation Thermal conductivity at 0 °C

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𝑘0 (1 + 𝑐𝑇) 𝑘0

𝑘(𝑇) =

0.4 μW/m3lower crust 10 km

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Depth where𝐴(𝑧) is 1/e of its surface value

0.02 μW/m3lithospheric mantle

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−𝑧

𝐴(𝑧) = 𝐴0 𝑒 ℎ𝑟

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T0

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and heat production) within the crust and lithospheric mantle.

2.5 W/m/K

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Coefficient of thermal conductivity 0.001 °C−1 𝑐 Table 2. Input parameters used to solve the earth thermal conductivity equation (Eq. 5, 6 and 7) to estimate the paleo-lithospheric thickness.

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The southern Patagonian Andes, to the west of the study region, were affected by thrusting and crustal-lithospheric thickening during most of the Cretaceous to middle Miocene (see Ghiglione et. al., 2016). These regions, consequently, cannot be considered thermally stable for our paleo-heat flow studies. However, eastward, from the foreland to passive Atlantic platform areas (cratonic zones, see Fig. 1), the region can be considered in thermal equilibrium, under a steady state condition (cf. Rodriguez and Littke, 2001; Sachse et al., 2016). Rodriguez and Littke (2001) and Sachse et al. (2016) estimated the paleo-heat flow history of the Golfo de San Jorge and Austral basins (see location in Fig.2a), respectively; using vitrinite reflectance

analysis on nine hydrocarbon boreholes, 25 samples each. Their studies have shown a constant heat flow value of ~50 mW/m2 for approximately thirty millions of years, since the Eocene (50 Ma) until Miocene (20 Ma), when the asthenospheric window started to develop under southernmost Patagonia (Fig. 1 and Fig 2a). These values were used in Equations 5-7 to calculate paleo-lithospheric thicknesses, following the Artemieva and Mooney (2001) approach. The thermal parameters and thickness of the crust used in calculations are shown in Table 2 (see also Ávila and Dávila, 2018 for further details).

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We calculated 𝑉𝑖 solving the second term of Eq. 4 using the Present-day lithospheric

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thicknesses (Ávila and Dávila, 2018) as in the residual topography calculations, and paleolithospheric thicknesses estimated for the late Miocene. It is important to highlight that Hc

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would not have changed from late Miocene until Today across the foreland (Lagabrielle et al., 2004). As described previously (see Geological and Tectonic Setting chapter), OIB basalts are

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subhorizontal and cover the easternmost Andean thrust front and foreland areas (see Fig. 1),

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suggesting that no major crustal deformation events occurred across these regions during this

5. Results

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time lapse.

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5.1 Residual topography

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Figure 4 Residual topography map of southern Patagonia in meters (m) Negative values (blues) evidence total topographies larger than isostatic, suggesting the presence of sublithosperic downwelling forces to match the excess of topography. Positive values (reds) indicate the opposite, whereas zero values (or close to zero) evidence equilibrium (white), when no additional forces are required to account for the observed topography.

Figure 4 shows our residual topography results derived from Present-day lithospheric

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model described above. It illustrates, from west to east (over the slab window region), positive residual topography values (red ~500 m) along the Andean belt, compensation in the

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Patagonian foreland (whites, close to zero) and negative values (blues ~200 m) eastward, along the Atlantic coast. Out of the slab window, the northern Patagonia shows a modern residual topography mostly null (i.e., isostatic equilibrium) to positive (reds <1000 m). From our results we can assert that although the residual topography model evidences

dynamic uplift along the Andean belt, they are null or close to zero in the foreland and zero to negative toward the Atlantic coast areas. Our residual topography calculation contrast substantially with previous dynamic topography models (e.g., Dávila et al., 2018; Guillaume

et al., 2009), which suggested ~1 km of dynamic uplift or positive residual topography, associated with the subduction and northward migration of the Chilean ridge and formation of a slab windows.

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5.2 OM mapping and estimations of uplift rates (𝑉𝑠 )

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Figure 5. Topographic reconstruction of the late Oligocene - early Miocene (OM) marine level. a) The 3D plot illustrates the OM level altitudes (black dots), planar tendency (PT) and predictions bounds (90%) (PB). Circles show areas out of the prediction bounds. b) Elevation map in meters (m) of the OM marine level Today. Dashed lines represent forebulge axes (<5 m height, see text for details) modelled with effective elastic thicknesses of 30 and 10 km (see Dávila et al., 2019). Solid black lines show the Andean thrust front location (easternmost thrusts).

Figure 5 illustrates the topographic reconstruction of the OM marine level, using a planar

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trend from 4078 data (extracted from a geological and topographic maps compilation, see methodology). The coefficient of determination shows an acceptable fit (r2 = 0.74, see Fig. 5a).

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It is important to notice that the only points that fall out of the 90% prediction bounds (Fig 5a) are those collected in the Andean belt and Macizo Del Deseado region (Fig. 1, Fig. 5a). These

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regions, in contrast to the remaining foreland, experienced crustal deformation after the Patagonian Miocene marine incursion (Folguera et al., 2018). If we disregard these “anomalous” data, the fitting is significantly better (r2 ~83%). In Figure 5b, we show the location of areas affected by thrusting (frontal thrust belt) and flexural deformation (two modelled forebulge axes, amplitude ~5 m, using effective elastic thicknesses of 30 and 10 km, respectively, cf. Dávila et al., 2019). Figure 5b also shows a digital elevation model of the OM marine level. The highest altitudes develop along the western side of the map (~1100 m asl to

the north, and 800 m asl to the south). There is also an elevation reduction trend from west to east, across the foreland area, with values ~200 m asl near the Atlantic coast. The OM horizons are not exposed in the southeastern most study region. In fact, seismic interpretations) (e.g., Ghiglione et al., 2019) and borehole data (see Dávila et al., 2019) in this region show the OM tops at tens of meters below surface (Fig. 5b).

5.3 Paleo-lithosphere and of uplift rate estimations (𝑉𝑖 and 𝑉𝑑 )

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From paleo-heat flow data (Rodriguez and Littke, 2001; Sachse et al., 2016, see location

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of boreholes used in these calculations in Fig 2a), we calculated the lithospheric thicknesses for the Miocene (before the slab window formation), in the thermally-equilibrated Patagonian

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foreland region. Values, depending on the heat flow used, ranged between 69 and 75 km (see Fig. 6a). Isostatic topography calculations (Tiso) suggests a surface near sea or below level for

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this lapse of time (~200 m asl Fig. 6a). This agrees with the occurrence of the OM level across

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the foreland (Cuitiño et al., 2016; Encinas et al., 2018). Our results show a major isostatic contribution (and minor or no dynamic topography) during the Miocene, in agreement with

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recent basin studies (Dávila et al., 2019).

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Figure 6. Surface uplift rates considering lithospheric thinning (𝑽𝒊 ) and comparison with observational topographic changes (𝑽𝒔 ) during the slab window formation. a) Right column shows the lithospheric thicknesses that result from solving the thermal conductivity equation (using paleo-heat flow data, thermal parameters, and crustal thickness of 30 km, see main text and Table 1 for details) for the Miocene (i.e., before the slab window). Right column shows the Present-day seismic and thermal lithospheric thicknesses (Ávila and Dávila, 2018; Robertson Maurice et al., 2003 respectively). SL: Sea level. b) Profile (west-east) from the thrust front (black dot) across the foreland comparing the 𝑽𝒊 results (black solid line) and OM marine level reconstruction (𝑽𝒔 , orange line). Notice the good fitting between 𝑽𝒊 and 𝑽𝒔 .

With the end-member values of lithospheric thicknesses (i.e., a Miocene paleolithosphere of 69-75 km and a modern lithosphere of 60-45 km), we estimated the vertical

motion of Earth surface in southern Patagonia. Figure 6b illustrates and compares 𝑉𝑖 results with the OM marine level reconstruction (𝑉𝑠 ) from the thrust front across the foreland (from west-east). Both curves 𝑉𝑖 and 𝑉𝑠 show a clear decreasing tendency from west to east, from ~1 km, in the westernmost foreland areas, to 0.5 in the Atlantic coastal margin. From the Equation 4, we can also estimate 𝑉𝑑 values. They are negligible across the foreland (100-200 m), indicating minor dynamic changes before and after the slab windows formation.

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6. Discussion

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6.1 Long-wavelength plateau uplift in southern Patagonia foreland

The Present-day elevations of the OM strata (see Fig. 5) might be explained by the

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interaction of (1) thrusting (mainly along the southern Andean belt) overlapped by (2) subcrustal forces (dynamic uplift and lithospheric mantle changes). However, no evidences of

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crustal tectonic activity have been reported in the extra-Andean foreland region (see Folguera

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et al., 2018). A likely explanation could be flexural upwarping (forebulge uplift). Dávila et al (2019) calculated the spatial and temporal flexural effects driven by tectonic crustal loading

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across the southern Patagonian foreland. This study demonstrated, from the early-middle Miocene, a flexural wavelength of ~200 km and maximum bulge amplitude of 5 m at ~100 km from the thrust front (see Fig 3, and Fig. 5 in Dávila et al., 2019). Evidently, this upwarping

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area along the bulge is not enough to reproduce the Present-day altitude of the OM strata and surface uplift in the foreland. Subcrustal forces are required to fit the modern topography of the OM tops (see Fig. 5).

The OM elevations are, however, not homogenous along strike (Fig. 5). This level is

~300 m higher to the north respect to the south region (see Fig. 5). This was likely generated during the middle Miocene foreland basin subsidence evolution (see Dávila et al., 2019) occurred after the OM strata accumulation (i.e., during the sedimentation history of Santa Cruz

Formation and correlatives). Ghiglione et al. (2019) have shown a remarkable northward reduction of shortening and stratigraphic thicknesses during the early-middle Miocene (after sedimentation of the OM strata, i.e., during the accumulation of the Santa Cruz Fm and correlatives). Recent backstripping and flexural computations, using outcrop stratigraphic successions and hydrocarbon boreholes (see Dávila et al., 2019), suggest that this tectonic loading event could have modified the OM position along strike. This is likely the main reason the OM strata crops out above sea level to the north and are in subsurface (or close to surface)

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under sea level to the southernmost Patagonia (see Fig. 5).

6.2 The modern lithospheric state and plateau uplift in southern Patagonia

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The spatial and temporal matching between the CTJ northward shifting and the increase in elevation of southern Patagonia since the late Miocene suggests that the relationship between

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oceanic ridge collision and relief evolution is not casual. However, a question remains to be

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answered: Which is the driving mechanism that relates these processes? Thomson et al. (2010), for example, linked the exhumation history, elevated topographies and orogenic growth in the southern Patagonian Andes with upper crustal processes (deformation driven by ridge collision

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(Ramos, 2005) and variations of topography along-strike as a result of changes in the glacial erosion efficiency). Other works associated the modern elevation in southern Patagonia with

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dynamic upwelling generated by the slab window evolution (Dávila et al., 2018; Dávila and Lithgow-Bertelloni, 2013; Flament et al., 2015; Guillaume et al., 2010, 2009). Studies on the elevated marine terraces (approximately 180 m asl) along the southern Patagonian Atlantic coast (see Pedoja et al., 2011) support this interpretation. Georgieva et al. (2016), more recently, rejected a dynamic topography contributions based mainly on thermochronological modelling and structural observations near the CTJ area. According to these authors, the lowtemperature thermochronology in the Patagonian Andes can be fully reproduced with upper

crustal faulting and erosion. However, these latter works obviated the occurrence of surface uplift and moderated-elevation plateau formation across the southern Patagonian foreland (see Dávila and Lithgow-Bertelloni, 2013), where crustal deformation or deep glacial erosion were not major controls. In our work, in addition to reassert the occurrence of relief formation (Fig. 6b) without crustal deformation across the Patagonian foreland, we also minimized the influence of dynamic topography contributions. Our results show nulls or small residual topographies (positives and negatives), smaller than those shown by existing models (Dávila

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et al., 2018; Guillaume et al., 2009), and more similar to Flament et al. (2015) results (~300

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m). It suggests that the main driving mechanism would be associated with changes in the lithospheric mantle. In fact, as shown in recent works (Davies et al., 2019; Flament, 2019), it

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is likely that the connection among both isostatic and dynamic processes at shallow depth (base of the lithosphere) is causal (see Fig. 7) and dissociated with large mantle convections in the

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lower mantle.

Figure 7. Geodynamic scenario of Southern Patagonia illustrating the relationship between the reduction of lithospheric thickness and dynamic topography (modified from Flament, 2019), triggered by the presence of an asthenospheric window, and generated by the subduction of a seismic ridge. CTJ: Chilean Triple Junction. Red arrows are upwelling flows and blue arrows downwelling flows. Note the difference in wavelengths (lambda) of dynamic topography motorized by the movements of the lower and upper mantle and those products of lithospheric thickness changes.

6.3 What drove surface uplift and plateau formation in southern Patagonia? The vertical movement of Earth surface could be generated by changes of isostatic topography, dynamic topography or a combination of both. Our Figure 6 shows these changes since the arrival of the oceanic ridge and slab window formation. It is clear that topography within the foreland areas (i.e., far away from the tectonic controls: thrusting and loading) was mainly driven by subcrustal forces. Dynamic topography, in turn, does not seen to have

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strongly influenced on surface uplift (see above). A potential driver is the lithospheric mantle

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thinning. Ávila and Dávila (2018) shown abnormally high heat flow values in southern Patagonia, consistent with the slab window formation (Breitsprecher and Thorkelson, 2009).

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They concluded that hot buoyant asthenospheric mantle rises up the isotherm of ~1300°C and, consequently, the LAB would place shallower than in those areas out of the window. These

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observations agree with Hager and Clayton (1989) (taken from Flament et al., 2013), who

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shown that at shallow depths, dynamic and isostatic topography are essentially one and the same (see the dynamic topography kernels in Fig. 3 of Flament et al., 2013). Davies et al. (2019) and Flament (2019) shown that, at shallow lithospheric depth, changes in lithospheric

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thickness can trigger small convection cells (Fig 7), which can generate small dynamic topography (like the one shown by our residual topography computations and Flament et al.,

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2015 numerical modelling of mantle convection). Considering that thermal analysis on Miocene units (before slab window) evidences lower paleo-heat flows and, therefore, thicker lithospheres, and that the crustal thicknesses would have not changed remarkably since then, we can assume that most of the lithospheric thinning in Patagonian was conducted by lithospheric mantle reduction. The isostatic uplift that results from the lithospheric thickness changes (𝑉𝑖 ), since the arrival of the slab window until Today, shows a remarkable fitting with the reconstruction of the OM marine units uplift (𝑉𝑠 ), during the same lapse of time. This might

suggest a causative relationship. However, there is not a perfect matching. The dynamic uplift, 𝑉𝑑 , which results from the subtraction of 𝑉𝑖 and 𝑉𝑠 , is indeed small (<300 m), playing a minor role in Patagonian surface development. This result, again, is in agreement with Flament et al. (2015), who shown ~225 m of cumulative uplift since 30 Ma in southern Patagonia. This allows us to suggest that, although recent works have disregarded dynamic uplift along the Patagonian Andes, it might have affected the Cordilleran belt in these amounts. However, these magnitudes are too small to be detected by geological studies. Nevertheless, we can assert that

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lithospheric thinning driven by slab window formation, particularly of the lithospheric mantle,

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would have been enough to reproduce modern elevations and surface uplift from Miocene to Present day. This thickness reduction, in turn, might have driven small mantle cells and modest

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and localized dynamic uplift (see Fig. 7).

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7. Conclusion

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Residual topography and uplift rates calculations across the southern Patagonian foreland suggest that Present-day isostasy and lithospheric thinning were the major control on uplift and plateau formation since the slab window opening in the late Miocene to Today. The uplift of

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the OM marine units (𝑉𝑠 ) matches with the lithospheric thickness changes (𝑉𝑖 ). On the other hand, asthenospheric upwelling or downwelling (i.e., dynamic topography and changes of

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dynamic topography, 𝑉𝑑 ) would have been <300-200 m in the study region. These values contrast with those estimated using forward numerical models (e.g., Dávila et al., 2018; Guillaume et al., 2013, 2009). Both contributions derived from our analysis, isostatic and dynamics, should be extrapolated to the Andean regions, to the west, where crustal deformation and glacial erosion overlapped with them to contribute on the total (or observed) topography. Our work, in turn, reinforces the hypothesis on the influence of oceanic ridge subduction and topographic uplift by lithospheric thinning and that, at least in Patagonia, slab window

formation is responsible of only a few meters of dynamic uplift. This could be of particular interest for analyzing other areas affected by slab windows (e.g., Becker et al., 2014). We do not discard the idea that lithospheric thinning could have trigger small mantle convection cells and small dynamic uplift. Author Statement

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Pilar Avila: Conceptualization, Methodology, Software, Validation, Formal analysis, Investigation, Resources, Writing - Original Draft, Writing – Review, Editing, Visualization, Supervision, Project administration Federico M. Dávila: Conceptualization, Methodology, Software, Validation, Formal analysis, Investigation, Resources, Writing - Original Draft, Writing – Review, Editing, Visualization, Supervision, Project administration, Funding acquisition

ACKNOWLEDGEMENTS

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We declare no conflict of interest of any type.

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Conflict of Interest

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We thank Irina Artemieva, Nicolas Flament and Andres Folguera for their valuable and constructive suggestions and reviews. We thank Juan Manuel Alcacer for provided us crustal

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thickness dataset and Juan Carlos Afonso for fruitful discussions. We appreciate funding from FONCyT (PICT funding program), SECyT-UNC, CONICET, PUE-CICTERRA 2016, PICT-

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E 2014 and 2018.

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