Measuring dynamic topography in South America

CHAPTER

Measuring dynamic topography in South America

3

Federico M. Dávila, Pilar Ávila, Federico Martina, Horacio N. Canelo, Julieta C. Nóbile, Gilda Collo, Francisco Sánchez Nassif, Miguel Ezpeleta Earth Science Research Center, CONICET National Research Council & Natural, Physics and Exact Sciences Faculty, Cordoba National University, Córdoba, Argentina

1 ­Introduction The modern Earth topography is the result of complex interactions between climate, crustal tectonics, lithospheric mantle, and asthenosphere. The landscape and the evolution of the basins, as part of the topography, respond to these changes that, in turn, can be preserved in the geological record. This suggests that topography is consequently transitory and dynamics, strongly influenced by the previous immediate states (see Dávila and Lithgow-Bertelloni, 2015). Crustal changes have been traditionally considered the main control on Earth surface topography. An example are foreland basins (e.g., Beaumont, 1981; Jordan, 1981 to more recent syntheses like Allen and Allen, 2013, among many others), which, in fact, have taught us that the main driving subsidence mechanisms occur by orogenic shortening during thrust sheet stacking in the uppermost crust. This creates tectonic loads, which bends the Earth surface creating accommodation spaces, where sediments (derived from elevated source areas, under favorable climate conditions) accumulate and preserve. Subcrustal forces (e.g., lithospheric changes or dynamic topography, see Lithgow-Bertelloni and Richards, 1998; Bjørnerud et  al., 2002) are considered, in turn, redundant and complementary. Assuming this, as the crustal loads increase, the associated basins deepen, and the sedimentary records should reflect this. On the other hand, surface uplift along a Cordilleran and foreland thrust belt is frequently associated with the cratonward migration and widening of the orogenic wedge (see Davis et al., 1983; McClay, 2011), which is related, in turn, with changes of crustal thickness by shortening. This drives not only the rising of the Earth surface but also drainage basin reconfigurations. Other subcrustal or sublithospheric processes have been considered poorly influencing. They are taken into account only when decompensation is very large, like in the Altiplano-Puna plateau (see Garzione et al., 2017) or in particular settings where tectonic deformation is absent (like the African superswell, Gurnis et al., 2000). Do we observe such a close relationship between mountain building and foreland basin development? Are the crustal changes or orogenic root formation the main driving mechanisms on Earth surface dynamics? From recent geodynamic modeling (Dávila and Lithgow-Bertelloni, 2013, 2015; Eakin et al., 2014; Dávila et al., 2018), the “noes” rules. What if crustal thicknesses kept constant and topographic changes are still recorded or if we detect subsidence and basin formation but the Andean Tectonics. https://doi.org/10.1016/B978-0-12-816009-1.00003-4 © 2019 Elsevier Inc. All rights reserved.

35

36

Chapter 3  Measuring dynamic topography in South America

lithospheric and/or dynamic forces are instead more consistent with uplift and generation of elevated plateaus at the same time? South America, and its evolution after the Gondwana fragmentation, is an excellent laboratory to study topographic changes. The recurrent marine incursions (Gianni et al., 2018b), basin stratigraphy (see Dávila and Lithgow-Bertelloni, 2013) as well as drainage system changes (Rodríguez Tribaldos et al., 2017) are, among others, excellent “tape recorders” to evaluate the driving forces. Another advantage is that the extended foreland areas exposes basalts (which assist us to estimate paleolithospheric structures) of different ages (e.g., Martina et al., 2017) as well as marker beds (Hoke et al., 2014) that can be used to calculate paleoelevations and to compare with the Present-day topographies. Here we revisited the topographic problems of South America using a different viewpoint. We make use of the residual topography concept (see more details in Flament et al. (2013) and Dávila and Lithgow-Bertelloni (2015)), which deals with the comparison between the elevation derived from compensation states and observed elevations. There are different methods to estimate it. But mostly they are based on global elevation quantifications (by removing the mean continental elevation of 529 m, e.g., Gurnis et al., 2000; Flament et al., 2013) and isostatic and/or flexural elevation estimations (by considering a crustal or lithospheric isostatic elevation and loads, see more details in Dávila and LithgowBertelloni, 2015). Although results differ strongly among models (Fig. 1), depending on the variables (layers and densities) and considerations used (type of model), these studies have demonstrated that crustal deformation and, likely lithospheric mantle changes, are not enough to reproduce different geological features (Ávila and Dávila, 2017). The blues and reds in our Fig. 1 represent excesses and deficits of topography, respectively, which might represent areas where sublithospheric forces compensate the isostatic forces to reproduce modern topography. Is it possible to account for these blues and reds at lower scales, for example, in sedimentary or drainage basins or for ancient settings? In fact, this residual concepts can be applied to backstripping and flexural models traditionally used in basin analysis. This, consequently, would allow tackling indirectly with the sublithosheric controls on topography. If we can estimate paleoisostatic states and paleoelevations, we could also calculate paleoresiduals.

FIG. 1 Residual topographic models of South America based on (A) a global average elevation of 529 m (modified Flament et al., 2013), and isostatic elevations using different lithospheric models: (B) CRUST 1.0 and Priestley and McKenzie (2013), and (C) LITHOS 1.0. Black polygons show the position of the major cratons.

2 ­ Geodynamic and tectonic setting

37

In this contribution, we propose to use published data and examples along the Andes chain and proximal and distal foreland areas in order to evaluate the main contributions on the South American topography. Our objective is to incorporate geological observation to test geodynamic hypotheses. To do this, we attack the problem by analyzing regional and local residual topographic proxies to compare not only among them, but also with dynamic topography models. Three “representative” transects across the northern-central, central, and southern Andes (Amazonas, Precordillera-Pampas, and Patagonia areas, respectively) were studied to test our hypothesis. These regions are influenced by oceanic ridge collision (seismic and aseismic), where we expect to record large geodynamic influences on subsidence and/or uplifting.

2 ­Geodynamic and tectonic setting The South American plate has been divided into three major tectonic areas that from west to east are: the Andean belt, cratonic regions, and South Atlantic platform (see Cordani et al., 2000, among others) (Fig. 2). We include in this work, only for practical purposes, a transitional region that likely behaves different from the rest: the pericratonic South American plains (see Dávila et al., 2010), located between the distal forelands and craton platforms (Fig. 2, pericratonic forelands from now). The subsidence and uplifting driving mechanisms across these four regions have differed remarkably since the Gondwana fragmentation. While the Andean belts have subsided and risen mainly by upper crustal deformation, shortening and bending conducted by the thrust sheet dynamics (see Jordan et al., 2001), with minor and local episode of stretching (e.g., Dávila and Astini, 2003), the cratonic and platform areas would be mainly affected by crustal stretching (see Heine et al., 2013), with minor and local shortening events. The subcrustal controls on topography also vary among regions. While the Andes are mainly affected by the subducting slab dynamics, through changes in the mantle wedge impacting on the lithospheric mantle (e.g., Fromm et al., 2004; Bishop et al., 2018), the cratonic to passive margins are mostly controlled by up- or downwelling “plumes” (see Dávila and Lithgow-Bertelloni, 2013; Davila et  al., 2018; Shephard et  al., 2010). These processes might also vary along strike the continent. In fact, the Andes and foreland were subdivided into three major segments (see Gansser, 1973; Mpodozis and Ramos, 1989): northern, central, and southern Andes on the base of main driving mechanisms on orogenesis. An important consideration for our analysis, and reference framework, is at the South American pericratonic forelands would have been poorly affected by crustal shortening and/or stretching since the end of the SW Atlantic rifting and beginning of the Andean orogeny; and placed near sea level since (at least) from the Cretaceous (Marengo, 2006, Dávila et al., 2010; Malumian and Nanez, 2011; Eakin et al., 2014; Reinante et al., 2014; Rodríguez Tribaldos et al., 2017; Gianni et al., 2018a,b). Today, these areas are between 150 and 0 m above sea level and average crustal thicknesses are between 38 and 37 km (see Pasyanos et al., 2014) (see Fig. 2). Nevertheless, the lithospheric structure could have changed (see later). We also assume for our analyses that the western plate margin was affected by a nearly vertical subducting subduction of the Farallones plate in the Late Cretaceous-Paleogene, as evidenced by a relatively stationary volcanic arc along to the modern Chilean forearc to High Andes. After the fragmentation of Farallones, the Nazca plate becomes more orthogonal (Somoza, 1998) and shallower in the early Miocene when arc volcanic rock starts to crop out eastward and flat in late Miocene to Present, especially on those segments affected by ridge subduction (Kay and Mpodozis, 2002). However, we cannot ignore that short and different stages might have locally affected this general setting. For example, in the modern segment

38

Chapter 3  Measuring dynamic topography in South America

FIG. 2 (see figure caption on opposite page) (Continued)

2 ­ Geodynamic and tectonic setting

39

between 37°–48°S, a Cretaceous flat subduction and later paleogene slab window (See Ramos and Kay, 1992; Aragón et al., 2013). We also propose that after the main stage of Neogene basin sedimentation close to the Andean chain, several topographic changes occurred along the Andean foreland, promoting the development of the modern fluvial system configuration (Rodríguez Tribaldos et al., 2017) and main uplift along the Cordillera (Hoke et al., 2014). In this revision, we are particularly interested in those regions affected by oceanic ridge subduction (both aseismic and seismic): the Nazca, Juan Fernández, and Chile ridges at 5°N–12°S, 30°S, and 46°– 52°S: (1) the Amazonas foreland in the northern-Central Andes, (2) the Precordillera-Pampas transect in the southern-Central Andes, and (3) southern Patagonia in the Southern Andes. These oceanic tectonic features affect(ed) not only the slab buoyancy and thermal structure of the upper plate (e.g., Collo et al., 2017), and consequently the isostatic state (see Ávila and Dávila, 2017), but also on the dynamic forces within the asthenosphere (Dávila and Lithgow-Bertelloni, 2015). In these areas, therefore, we expect to find stronger subcrustal controls on Earth surface topography than on other segments. Now we describe these three transects very synthetically. For further details, refer to specific chapters of this book. The northern-Central Andes, the Acre-Solimoes-Amazonas foreland spans 2000 km along and across strike, between approximately 5°N and 12°S (Fig. 2), on a segment affected by slab flattening and attributed to the subduction of the buoyant aseismic Nazca ridge since late Miocene time (Eakin and Long, 2013). The region is clearly divided along strike by the Fitzcarrald Arch (Espurt et al., 2007; Bishop et al., 2018), a W-E topographic feature that break the foreland into two basins: the (a) Marañon-Ucayali-Acre to the north and (b) Madre de Dios to the south (Räsänen et al., 1990; Mathalone and Montoya, 1995; Latrubesse et al., 2010). Although the onset of synorogenic sedimentation in foredeep has been proposed Late Cretaceous to Paleogene (Campbell et al., 2001; Hermoza et al., 2005; Hoorn et al., 2010), the main and thickest Andean successions are Miocene, locally near 2000 m thick. Further east two important depocenters develop in pericratonic to cratonic areas, separated by two highs (Carauari and Purus), the Solimoes and Amazonas basins, which also record approximately 300–500 m of Cenozoic strata (Caputo and Soares, 2016; Heimdal et al., 2018). Along the Atlantic margin, high accumulation rates were reported in the marine megafans located in the Amazon River mouth (Figueiredo et  al., 2009). The foreland basin formation correlates with an important FIG. 2, CONT’D Geodynamic setting of South America (overlying an ETOPO1 image) showing different geographic, tectonic, and geodynamic features described in the text. Note the location of the three transects (gray lines) facing the subduction of oceanic ridges (NR, JFR, and CR). The Northern (NA), Central (CA), and Southern Andes (SA). Dashed white line indicates the location of the pericratonic foreland, between the flexural-driven foreland and cratons. Yellow dashed line indicates the extension of the Peruvian flat-slab segment. Main features in the Peruvian flat-slab region (red labels) are: M, Marañon; U, Ucayali; FA, Fitzcarrald arc; MA, Madre de Dios. Magenta dashed line indicates the extension of the Chilean or Pampean flat-slab segment. Main features in this region (red label) are: CC, Coastal Cordillera; PFC, Principal and Frontal Cordillera; P, Precordillera; B, Bermejo; SP, Sierras Pampeanas. Red dashed line indicates the approximate extension of the slab window on the Patagonia region of Chile and Argentina. Red dot: Triple Junction (TJ) develop by the joining of the Nazca, Antarctic, and South American plates. MA: Magallanes or Austral basin. Black arrows show the direction of the Nazca and Atlantic plates movement relative to South America plate and black dots are the borehole locations modeled using backstripping (see text for details).

40

Chapter 3  Measuring dynamic topography in South America

e­ xhumation in the Peruvian Andes, recently constrained by zircon and apatite (U-Th)/He ages between the Oligocene to Miocene (Scherrenberg et al., 2016). It is important to mention that the foreland strata in the northern-Central Andes are underlain by Cretaceous successions, which are locally covered by transitional and/or marine beds (Alter do Chao Formation, Rossetti and Netto, 2006) that have been regionally correlated along strike. The Neogene foreland basin geometry is not a typical wedge foredeep but rather regular across strike (Dávila and Lithgow-Bertelloni, 2013). While Espurt et al. (2007) documented 2000–1500 m of synorogenic strata to the west, toward the proximal foredeep, Latrubesse et al. (2010) reported comparable strata thicknesses in the Acre basin, 400 km away from the easternmost Andean load at these latitudes as well as Caputo and Soares (2016) in the Amazonas. Flexural calculations have not been able to reproduce this geometry (see Eakin et al., 2014) and additional loads were required to accommodate the Miocene synorogenic strata (Dávila and Lithgow-Bertelloni, 2013). Recent estimates of shortening along the Peruvian Andes (Scherrenberg et al., 2014, 2016), however, might assist to improve our previous analysis and revisit interpretations about deficits or excesses in the creation of topography in the Amazon sections (see Gotberg et al., 2010; Scherrenberg et al., 2014). These newest balanced crosssections allowed explaining ~130 km (38%) of Cenozoic Andean deformation. Although this shortening is rather similar to our previous work (Dávila and Lithgow-Bertelloni, 2013; Eakin et al., 2014), the load distribution differs. Exhumation and uplift across the Peruvian flat-subduction segment, related to the formation of the Fitzcarrald Arch, have been associated with the Nazca ridge collision and dynamic uplift given that across the pericratonic foreland tectonics was meager (see Regard et al., 2009; Eakin et al., 2014; Dávila and Lithgow-Bertelloni, 2015; Bishop et al., 2018). However, these interpretations are based completely on modeling. A few contributions (Regard et al., 2009) have performed direct observations of this morphostructural feature. This is because active erosion is likely much larger than rock uplift and, consequently, surface uplift is close to null. This was clearly shown by numerical simulations (see Braun et al., 2013). Boreholes across the Brazilian pericratonic foreland (Acres, Solimoes, and Amazonas) extracted from Caputo and Soares (2016) and Heimdal et  al. (2018), allowed us to test previous interpretations on nonsedimentary subsidence and dynamic topography (Shephard et al., 2010; Dávila and Lithgow-Bertelloni, 2013; Eakin et al., 2014). In the south-Central Andes, between 27° and 34°S, the Andean tectonics has also been mainly governed by changes in the slab angle, from shallow to flat since the middle Miocene as a result of the subduction of a buoyant aseismic ridge, known as Juan Fernández (Yáñez et al., 2001; Kay and Mpodozis, 2002). Before this stage, from (at least) the Cretaceous to early Miocene, subduction would have been nearly continuous along this segment, dominated by plugging subduction (subvertical to normal dipping, i.e., >30° to the E), as suggested by the position of coeval volcanic-arc rocks in Chile (Mpodozis and Ramos, 1989; Schellart, 2017), and oblique (Somoza, 1998). Two major tectonic features are recognized across this transect; from west to east they are: (1) the Cordilleran belts (Coastal, Principal, and Frontal Cordilleras), to the west, and (2) the foreland areas to the east (proximal: Argentine Precordillera, and distal pericratonic: Sierras Pampeanas and Chaco or Pampas Plains) (Fig. 2). The westernmost Cordilleran belts, mostly developed in Chile, were tectonically active from the Cretaceous to early Cenozoic, as a consequence of the westward displacement of the South American plate. This belt suffered poor deformation and uplift from the middle Miocene onwards (see Hoke et al., 2014 and references therein). It would have behaved passively during the structural stacking history of the Argentine Precordillera thrust belt located forelandward. The Precordillera, in contrast, accommodated around 100 km of crustal shortening from the middle Miocene to Today (Jordan et al., 1993, 2001). The shortened areas would have been the main load and explanation to reproduce the large foredeep

2 ­ Geodynamic and tectonic setting

41

subsidence (Charrier et al., 2015; Horton, 2018; Mackaman-Lofland et al., n.d.) as well as the flooding of Tertiary marine tongues in pericratonic foreland areas (Ramos, 1999). Today, however, we know the flexural wavelength is not large enough to explain the subsidence or the displacement of the Earth surface below sea level (see Dávila et al., 2010). The Miocene foreland at these latitudes, known as the Bermejo basin, is 300 km wide and covers the Argentina Precordillera and Sierras Pampeanas regions. It shows a typical and very thick coarseningupward sequence (cf. Jordan et al., 1993, 2001), locally >10 km thick, that is interrupted upsection with the development of the broken foreland coarse strata (Dávila and Astini, 2007). Jordan et al. (1993, 2001), Cardozo and Jordan (2001), Milana et al. (2003), and Dávila et al. (2007, 2010) summarized the tectonostratigraphic evolution across strike. The maximum alluvial thicknesses (>10 km) were reported close to the thrust front, in the proximal foredeep in the Precordillera belt, whereas the distal pericratonic areas in the Sierras Pampeanas and Pampas, a hundred meters of strata have been reported (Dávila et al., 2007, Marengo, 2006; Dávila et al., 2010, 2018; Giménez et al., 2011). After the classical foreland sedimentation, the distal foreland areas were involved in deformation, developing a broken foreland basin setting from the late Miocene to Present (see Jordan and Allmendinger, 1986). Here, the intermontane areas (among basement thrusts) have been poorly affected by deformation but exhumed and dissected by Quaternary fluvial systems (Ezpeleta et  al., 2006). Dávila and Lithgow-Bertelloni (2013, 2015) proposed that these basins show long-wavelength upwarps, which might be associated with the arrival of slab flattening to the region (dynamic uplift, Dávila and Lithgow-Bertelloni, 2015). The absolute altitudes locally >1 km above sea level, and the 100-m relief of flat mesas affecting Miocene strata (Ezpeleta et al., 2006; Dávila et al., 2007), with poor tectonic activity, suggest a change from net subsidence to uplift (Dávila and Lithgow-Bertelloni, 2013). The Andean foreland strata are regionally underlain by upper Cretaceous rift beds, which record the first and oldest sedimentation stage after the Gondwana fragmentation. While Cretaceous is relatively thin (a few hundred meters) in the Precordillera (Ciccioli et al., 2005), large thicknesses were reported toward the pericratonic foreland, in the eastern Sierras Pampeanas and Chaco or Pampas plain (see Schmidt et al., 1995; Webster et al., 2004). Two marker beds stand out from the Cretaceous sedimentary pile, the “Serra Geral” basalts (140 and 70 Ma, see Lucassen et al., 2007; Lagorio, 2008) and the Maastrichtian (Reinante et al., 2014) to Paleogene (Marengo, 2015) marine strata, which developed mostly in subsurface. We use these two markers to analyze the Cretaceous-Paleogene paleotopography (see later) and account for dynamic forces previous to the Andean orogeny. While dynamic subsidence associated with subducting plunging slabs has been invoked to facilitate the marine incursions within pericratonic areas since the Gondwana fragmentation (MaastrichtianPaleogene and Paranaense flooding, Malumian and Nanez, 2011; Ruskin et al., 2011; Reinante et al., 2014, and references therein) and the accommodation of tens of kilometers of alluvial records in the Miocene, slab flattening during Mio-Pliocene to Present could explain the regional exhumation and uplift history (Dávila and Carter, 2013) across the Sierras Pampeanas (Dávila and Lithgow-Bertelloni, 2013, 2015) and subsidence toward the pericratonic foreland (Dávila et al., 2010, 2018). The effect of mantle dragging forces in the Chaco or Pampa Plains at the leading edge of the flat slab would have persisted until today, as evidenced by Quaternary stratigraphy and modern geomorphology (Dávila et al., 2010). In this contribution, we revisit this transect in order to test our hypotheses. We carry out a preliminary and novel approach based on the evaluation of the chemical nature of pre-Andean basaltic rocks (primitive melts), related to the paleolithospheric column. This will allow understanding the changes of paleotopography during the Gondwana breakup times. We analyzed with this purpose the Cretaceous basalts exposed in the Sierras Pampeanas (Lucassen et  al., 2007), following some methodological

42

Chapter 3  Measuring dynamic topography in South America

recommendations. In addition, we complement our previous subsidence studies by performing a backstripping analysis in boreholes located hundreds of kilometers east from the nearest tectonic load within the Chaco or Pampas Plain. We also take the chance to revisit previous flexural models (Cardozo and Jordan, 2001; Dávila et al., 2007; Fosdick et al., 2014) and structural reconstructions (Allmendinger and Judge, 2014) in order to quantify intermediate loading stages and flexural subsidence between the middle Miocene to Present (see Sánchez et al., 2019). This result was also correlated with a subsidence curve computed from a borehole located in the Bermejo foredeep. This might give some light about the arrival of the dynamic subsidence wave to the foreland previous to the flat-slab stage, apparently dominated by dynamic uplift (Dávila and Lithgow-Bertelloni, 2015). The foreland uplift history, occurred after the arrival of flat subduction in the late Miocene-Pliocene, has been however more difficult to evaluate from a dynamic topography perspective. The best places to do it are the most distal foreland areas, given that tectonics was scarce, but local relieves <500 m discourage paleoelevation studies given that the proxy errors are close or larger than relief. Therefore, we analyzed the regions westward across the flat-slab segment to re-evaluate the regional uplift history during flat subduction, revisiting recent paleoelevation approaches along the Andes at these latitudes, where vertical displacements would not be linked to tectonics (Hoke et al., 2014). The southern Patagonian Andes, between 45° and 55°S, in contrast to the South American segments described earlier, were affected by the subduction of the seismic Chile ridge, where a Triple Junction (TJ) develop by the joining of the Nazca, Antarctic, and South American plates (Fig. 2). Northward of the TJ, the subduction of Nazca plate has been “normal” since the Miocene, in terms of dip angle and extent within the mantle, whereas to the south the Antarctic plate subduction has been negligible, only 100 km beneath South America plate (Breitsprecher and Thorkelson, 2009). This has been supported by seismic studies along strike (Heintz et al., 2005, and our Fig. 3) and led several authors to propose a foreland evolution and topography (see Dávila et al., 2018) dominated by the formation of a slab window (Ramos and Kay, 1992; Kay et al., 1993; Gorring et al., 1997, Gorring and Kay, 2001; Breitsprecher and Thorkelson, 2009; Russo et al., 2010) (see Fig. 2).

FIG. 3 Diagram showing the along-strike distribution of seismicity with respect to depth from 55° to 20°S latitude. Notice the deepest events in Patagonia are <100 km, supporting a very shallow subducting slab (INPRES database, www.inpres.gob.ar).

2 ­ Geodynamic and tectonic setting

43

The southern Patagonia geology consists of a low-elevation Cordilleran thrust belt (elevation under 3000 m above sea level, Ramos, 2005) and a typical foreland basin system, known as the Magallanes (or Austral) Basin (see Biddle et al., 1986; Ramos, 1989; Ghiglione et al., 2010, 2016) (Fig. 2). The main foreland sedimentation occurred at these latitudes between the Late Cretaceous (c.101 Ma) and Early Miocene (Fosdick et al., 2011; Georgieva et al., 2016) and the uppermost and youngest foreland strata are ~14 Ma (Santa Cruz Fm., Blisniuk et al., 2005; Cuitiño et al., 2016). After Miocene sedimentation and the subduction of oceanic seismic ridge, the south Patagonian foreland started to uplift (Guillaume et al., 2009, 2013; Dávila et al., 2018). This is supported by a 6-Myr protracted unconformity (from 14 to 8 Ma), covered by the gravel sheets of the “Rodados Patagónicos” (<10 m thick, Parras et al., 2008) and development of fluvial terraces (Guillaume et al., 2009) carving the flat and elevated (>500 m above sea level) plateau. These gravel beds evidence fluvial bypassing to the passive Atlantic margin (cf. Ghiglione et al., 2016). Extensive tholeiitic plateau lavas <12 Ma (Gorring et al., 1997) lay on the Patagonian plateau, which postdate the latest foreland sedimentation (i.e., ~14 Ma). Basalts are not deformed and slightly tilted to the E (Lagabrielle et al., 2004). This implies that the main shortening and crustal tectonics would have ceased in south Patagonia at ~13 Ma, coeval with the plateau formation and northward migration of the seismic Chile ridge (Breitsprecher and Thorkelson, 2009). Our ongoing geochemical studies (Martina pers. comm.) will assist us to understand the lithospheric state during Miocene basalt formation, which might give us new tools to understand the topographic changes during the slab window formation. From these two main stages arise a subsidence period followed by an uplift history. The orogenic shortening in south Patagonia varies between 45 and 22 km since the latest Cretaceous, representing an accumulated deformation of ~10%–20% (Ramos, 1989; Fosdick et  al., 2011; Ghiglione et  al., 2014). Flexural analyses based on diverse load reconstructions (see Ghiglione et al., 2010; Fosdick et al., 2014), nevertheless, have not been able to reproduce the basin geometries (see discussion in Dávila and LithgowBertelloni (2013) and Dávila et al. (2018)). Additional loads are required to accommodate from 8 km (Ramos, 1989) to 5 km (Ghiglione et al., 2010) of strata distributed across the Patagonian foreland (Dávila and Lithgow-Bertelloni, 2013). Guillaume et al. (2009) and more recently Dávila and Lithgow-Bertelloni (2013) suggested that dynamic subsidence might have contributed with ~1–2 km of accommodation space during the early Miocene, generated during a plunging subduction stage prior to the arrival of the Chile ridge. A similar geodynamic scenario might also assist to explain the marine and deep Cretaceous depocenters reported by Biddle et al. (1986), which would require (if the subsidence is fully flexural-­ controlled) of a very large loading episode. Fosdick et al. (2014), in order to solve this problem, proposed a lithospheric attenuation (i.e., extremely low effective elastic thicknesses Te). Nevertheless, such a thinned lithosphere is not consistent with below-sea-level topographic surfaces. Lighter lithospheres (as shown below) produce elevations above sea level (e.g., the rising of the Andean plateau by delamination). In fact, recent studies have suggested during the Cretaceous a large-scale flat subduction stage (Gianni et al., 2018a,b) in order to explain, among other features, the shifting of subsidence toward the distal foreland that cannot be accounted by flexural models. In this contribution, we revisited the most recent structural reconstructions (see Ghiglione et al., 2014) in order to retouch and polish our previous flexural models. We also revise the Cretaceous reconstructions using more realistic effective elastic thicknesses. We expect under an attenuated lithosphere, to reproduce the foreland subsidence amplitude but not the sea levels (see Fosdick et al., 2014; Gianni et al., 2018b), which require a lithosphere explanation. From the Late Miocene to the present, a very incipient subduction would have affected the southernmost Patagonian margin. This is because the Antarctic slab was detected at only <100 km depth (Heintz et al., 2005). Such a geodynamic setting could favor (but not necessarily) the formation of upwelling mantle cells, which drive dynamic uplifting (Guillaume et al., 2009, 2013; Dávila et al., 2018).

44

Chapter 3  Measuring dynamic topography in South America

Two models have been proposed in Patagonia, which account for dynamic elevations (by Guillaume et al., 2009, 2013 and by Dávila et al., 2018), both based on paleogeographic reconstructions and northward migration of the Chile ridge. This uplift episode was also measured along the coastal Atlantic cliffs (Pedoja et al., 2011). These works documented an abrupt change in the paleoshoreline elevation at approximately the TJ latitudes. In this work, we compare and contrast these ideas with some preliminary residual topography results (Ávila and Dávila, 2017; Fig. 2), that are part of our studies in south Patagonia (Fig. 4).

3 ­Theory, methods, and data In this contribution, we revise the dynamic topography hypotheses of South America (see Dávila and Lithgow-Bertelloni, 2013, 2015; Dávila et al., 2018) by revisiting recently published geological and geophysical data (and not considered in our previous contributions), as well as incorporating a new and preliminary approach. To do this, we broaden the use of the residual topography concept to past geological times. Residual topography (e.g., Flament et al., 2013) in geodynamics is frequently associated with equilibrium analyses, where theoretical Earth elevations are compared with the observational elevations extracted, for example, from DEMs (e.g., SRTM 90-m). Residuals might also derive from comparisons between diverse topographic marker levels. Among others, the estimation of a basin substrate level, after removing sedimentary loads using backstripping analysis or considering tectonic load using flexural models (e.g., Cardozo and Jordan, 2001; Spasojević et al., 2008). Both modeled surfaces might find excesses or deficits of topography when compared with corrected observations, like stratigraphic columns (1D) or isopach maps (2D and 3D). While the latter cases are excellent proxies for negative topographic residuals, uplifting analyses are more complicated. Tectonically stable areas, like the pericratonic belts of South America in Amazonas, Pampas, and Patagonia (see Fig.  2), are however excellent places to look for residuals, where Cenozoic strata and marine beds were reported. For ancient settings, we look for those areas where the crust would have no changed from tens of millions years ago and the dynamic components can be estimated using subsidence approaches (e.g., see more details later). This allows us to estimate sensible data for present states and extrapolate them to calculate paleoisostatic elevations and, by discarding, estimate dynamic influences. These ambiguous data (and very sensible) for residual studies are the density contrasts between the diverse lithospheric layers and a supposedly homogeneous asthenosphere.

3.1 ­Global to continental residual topography The local elevation of the surface (Tt) can be described as a combination of an isostatic and a dynamic component. Tt = Ti + Ti

(1)

The residual topography results from the difference between the observed elevations (total topography) and the isostatic topography, which results from isostatic equilibrium (Flament et  al., 2013; Dávila and Lithgow-Bertelloni, 2013, 2015). This residue (difference) is usually associated with the deep component.

45

(A–C) Instantaneous flow models from Dávila and Lithgow-Bertelloni (2015), showing: (A) dynamic topography using a plunging subduction (30°E), (B) dynamic topography using flat subduction, and (C) the change of dynamic topography between (A and B). Notice that this model evidences a neutral change of dynamic topography in most South America, except in the flat subduction segments. (D and E) Instantaneous dynamic topography models from Patagonia in the southern Andes (after Dávila et al., 2018), depicting the results considering a subduction kinematics since the 12 Ma, when the Chile ridge started to affected the region from south to north, and (F) the change of dynamic topography between (D and E). Notice that Patagonia evidences a positive change, suggesting dynamic uplift.

3 ­ Theory, methods, and data

FIG. 4

46

Chapter 3  Measuring dynamic topography in South America

Lachenbruch and Morgan (1990) have shown that the isostatic equilibrium can be considered as a sum of contributions from the buoyancy of the crust (Tc) and the lithospheric mantle (Tml).

ε = a (Tc + Tml − H 0 ) where

 ρa  a =  ρa − ρw   1

ε <0 ε >0

, Tml =

(2)

ρ a − ρ ml ρ − ρc Hc ; H ml , and Tc = a ρa ρa

The elevation of a lithospheric column ε with respect to a reference elevation (H0), usually taken as the average midoceanic ridge elevation or any reference datum (2.4 km in Lachenbruch and Morgan, 1990), is given by the lithosphere buoyancy, with components from the crustal layer with a thickness (Hc) and density (ρc) and the mantle lithosphere with a thickness (Hm) and density (ρm), floating over an asthenosphere of density (ρa). While Tc is always positive because the crust is lighter than the asthenosphere, Tml is negative because the mantle lithosphere is heavier than the asthenosphere (same composition but cooler). Therefore, a region with a thicker lithospheric mantle layer should reproduce lower surface elevations. The Hc and Hml data were taken from global sources such as CRUST 1.0 (Laske et  al., 2013), LITHOS 1.0 (Pasyanos et al., 2014), and Priestley and McKenzie (2013). We use a DEM (from the ETOPO1) model to represent the total topographic elevation (Tt).

3.2 ­Local to regional residual topographic analysis These studies use different proxies to estimate the residuals within a specific area. They are based on different models and proxies. We now describe the two main approaches used in this work.

3.2.1 ­From paleoelevations by estimating paleolithospheric thicknesses

This method can be divided into two parts, a geochemical analysis using dated basalts with available geochemistry and modeling to estimate paleolithospheres. With these data, we can use the approaches stated here to calculate a paleoisostatic elevation. Although basalt are punctual data, they “drag” information from large areas below and might represent a lithospheric state larger than the area where were collected. We analyzed samples from Sierras de Córdoba where the Cretaceous magmatism is well recorded and has a complete geochemical database. Basalt information comes from Kay and Ramos (1996), Lucassen et al. (2007), and Lagorio (2008). The lithospheric modeling is based on Plank and Forsyth (2016) petrology method, considering a crustal thickness of 37 km (cf. Lucassen et al., 2007 and references therein). The method requires that melts have been extracted from a peridotite mantle source. However, since most samples are not primitive basalts, they were initially corrected for fractional crystallization by adding olivine until composition was in equilibrium with typical mantle olivines (Fo > 89), using the PRIMELT3 MEGA (Herzberg and Asimow, 2015) software. Only samples with >9.5% MgO that followed the liquid line of descent for crystallizing primary magmas were used to avoid major effects of cpx fractionation and crustal assimilation. The basic idea is to estimate firstly the modern isostatic lithospheric topography (thicknesses and densities) to account for the modern altitudes along these plains. We choose the pericratonic plains given that they were poorly affected by tectonics for the several tens of millions of years (Lucassen et al., 2007 and references therein). Assuming we can extract dynamic subsidence from backstripping

3 ­ Theory, methods, and data

47

analysis across stable areas (with no tectonics) and the lithospheric thicknesses from different catalogs (e.g., LITHOS 1.0), crust and mantle densities would be the only variables to adjust. These can be then extrapolated to ancient examples in order to estimate, for discarding, the ancient dynamic influences during the Cretaceous, after the fragmentation of Gondwana. The dynamic components are compared with global dynamic topography calculations (see later) for the past setting.

3.2.2 ­From subsidence analysis

They are based on two acknowledged basin methods, the elastic flexural and backstripping modeling. For both, we use the numerical solutions and programs written by Nestor Cardozo (Flex2D and OSXBackStrip). For more theory and numerical solutions, readers can refer to Cardozo and Jordan (2001) and Allen and Allen (2013). For our (a) flexural computations, the load shapes were drawn and discretized in rectangles of 10 km width from balanced cross section from three regional transects from Peru, Precordillera, and southern Patagonia (Gotberg et  al., 2010; Scherrenberg et  al., 2014; Allmendinger and Judge, 2014; Sánchez et al., 2019; Ghiglione et al., 2014), northern-Central, Southern-Central, and Southern Andes, respectively. These transects synthesize the tectonic and geodynamic behaviors from north to south. We used effective elastic thicknesses (Te) between 20 and 40 km (cf. Tassara et al., 2007), depending on distance from the trench (lower in the hinterland and higher cratonward), an infinite plate model, and Young and Poisson values of 70 GPa and 0.25, respectively (see Turcotte and Schubert, 2014 for further details). In the Precordillera, our model is quite different from our previous ones. There we reproduced intermediate reconstructions, in four steps between the middle Miocene to Present at 14, 12, 6, and 0 Ma. To build the loads, we use a different approach (see details in Sánchez et al., 2019), which considered a wedge geometry during each time step that results from the restitution of the thrust sheets and erosion through dipping surface. This surface preserves the observed stratigraphic records (exposed today in a cross section) but considering a subcritical-critical taper wedge geometry (cf. Dahlen, 1990). The (b) backstripping analysis assumes that water and sediment loads are compensated locally by the displaced weight of a column of the weak mantle (the asthenosphere), and that the porosity of the sediments decreases exponentially with depth as the sediments are buried (see Allen and Allen, 2013). This analysis consists of removing the effects of sediment compaction, and water and sediment loading. For this chapter, we focused in those areas poorly or no affected by crustal tectonics. Consequently, the residues estimated from this allowed us estimating the nonsedimentary components, assigned to sublithospheric forces or negative dynamic topography (subsidence). The input data for each model are thicknesses, ages, lithofacies, petrophysical properties, and stratigraphic relationships. Stratigraphic and sedimentological data were extracted from boreholes of the Austral and Golfo de San Jorge basins for Patagonia, Chaco or Pampas in the south-Central Andes, and from the Solimoes and Amazon basins in the northern-Central Andes (Fig. 2). Petrophysical properties (porosity and grain densities) were taken from Allen and Allen (2013).

3.2.3 ­From paleoelevation changes using paleoaltimeter proxies

We have not performed this type of paleoaltitude analyses in this work. We use data already published (Hoke et al., 2014) to understand the topographic evolution along the Argentine Precordillera to High Cordillera transects at approximately 30–35°S. For more methodological details, particularly on those used in this contribution (stable isotopic proxies), readers might refer to Kohn (2007). In thrust areas, the residuals are tricky to compute. We should know the displacements, structure geometries, and kinematics. With these, the changes of elevation of a certain block of rock might be compared

48

Chapter 3  Measuring dynamic topography in South America

with paleoelevation proxies, for example, the altitude estimated using stable isotopic composition. An alternative is to take advantage of those areas that we know, from tectonic studies, that thrusting and upper-crustal deformation have been quiet to null in the lapse of the paleoelevation estimations. However, those regions are today at <500 m above sea level, and errors overlap with modern elevation. Here the paleoelevation changes are revisited and reanalyzed using the residual topographies that result from the present-day state of the art of the crust and lithosphere at these latitudes. We also compare the results with published dynamic topographic changes (Dávila et  al., 2018 and GPlate Portal, see red and blue curves in Fig. 6B).

3.2.4 ­From longitudinal profile river modeling

These studies are based on the Pritchard et al. (2009) approach (recently sophisticated with different algorithms, e.g., Rudge et al., 2015), where uplift rates are calculated as a function of space and time from large sets of longitudinal river profiles. This strategy assumes that the shape of a river profile is controlled by the history of uplift rate and moderated by the erosional process. This analysis assumes that upstream drainage area is invariant. More recently, continental-scale drainage networks are supposed to record regional uplift patterns, with a continent-wide spatial coverage. Simultaneous inverse modeling of a large inventory of longitudinal rivers profiles allows estimating regional uplift. Likewise in the previous paleoelevation proxy analysis (2.c), we do not make calculations but rather we use recent large-scale models in South America (Rodríguez Tribaldos et al., 2017) to understand the uplift history and possible controls. We compare these results with ours derived from more local river modeling in Sierras Pampeanas (Nóbile, 2013; Canelo, 2015) and Patagonia (Ávila, 2015).

3.3 ­Dynamic topography We plotted in our Fig. 5 for different correlations the computations by Dávila and Lithgow-Bertelloni (2015), for South America, and Dávila et al. (2018), for Patagonia (see Fig. 5D and E). These computations are based on instantaneous viscous flow induced by a given prescribed density heterogeneity field by solving the conservation of mass and momentum equations for a Newtonian incompressible mantle in a spherically symmetric shell via propagator matrices (Hager and O'Connell, 1981). This model solves for stresses and velocities to spherical harmonic degree and order 50 and computes the dynamic topography as h = −Δρrr/g, where h is the total elevation, rr are the radial stresses induced by the flow, Δρ is the density contrast typical of the crust-air interface of 3200 kg/m3, and g is the gravitational acceleration. The backstripping curves (Fig. 6) as well as the Cretaceous dynamic topography of South America were also based on the Müller et al. (2018) work, compiled from the GPlatesPortal (see portal.gplates.org), which is a time-dependent flow adjoint inversion model (see Liu and Gurnis, 2008).

4 ­Results 4.1  ­Continental lithospheric scale analysis. Modern and ancient Fig.  1 shows our preliminary most recent results on residual topography at the scale of the whole South America. These computations, based on the CRUST1.0 and the Priestley and McKenzie (2013) lithospheric mantle model, and LITHOS1.0 data, together with crustal densities as function of crustal

4 ­  Results

49

FIG. 5 Flexural model results along the three transects, (A) the northern-Central Andes, (B) south-Central Andes, showing the (C) subsidence evolution from stratigraphic analysis and our flexural results (after Sánchez et al., 2019); and (D) Southern Andes. Modeling details and load reconstructions in text.

400

350

300

250

200

150

100

50

0

0 100

400

200

300

300

200

400 500

100

600 700

0

800

–100 –200

Depth (m)

Dynamic topography

450 500

Amazonas-Solimoes basins (northern Central Andes)

900 1000

300 200 150 100 50 0 –50 –100 –150 –200 –250 –300 –350

250

200

150

100

50

0

0 200 400 600 800 1000 1200 1400

Depth (m)

Dynamic topography

(A)

1600 1800

Chaco or Pampas basin (southern Central Andes)

2000

(B) 140

120

100

80

60

40

20

0

0

50

200

0

400

–50

600

–100

800

–150

1000

–200

1200

–250

Depth (m)

Dynamic topography

160 100

1400

–300

1600

–350

1800

Golfo San Jorge basin (Southern Andes)

–400

(C) 140

120

100

80

60

40

20

0

0 100

100

200 0

300 400

–100

500

–200

600 –300 –400

FIG. 6

160

Depth (m)

Dynamic topography

180 200

700

Austral basins (Southern Andes)

800

(D)

Decompaction analysis of a 1D Airy backstripping with exponential reduction of porosity (Allen and Allen, 2013) using the OSXBackstrip software (www.ux.uis.no/~nestor/work/programs.html) along the three transects in the (A) northern-Central (Solimoes and Amazon basins) (B) southern-Central (from the Bermejo basin (dash line) to Pampas basin (solid lines)) and (C) and (D) Southern Andes (Golfo San Jorge and Austral basins, respectively). Red and blue curves are the dynamic topography through time from Flament et al. (2015) and Müller et al. (2018).

4 ­  Results

51

thicknesses (cf. Zoback and Mooney, 2003), can be seen between 2700 and 2900 kg/m3. Lithospheric mantle density were calculated at 3280 kg/m3, by equilibrating the observed topography and modeled topography in the pericratonic foreland, where we know thicknesses of crust and mantle lithosphere as well as the dynamic component derived from backstripping (see later). Our preliminary studies show (see Fig. 1) three important features. First, both models differ remarkably in absolute values but show similar trends in some regions. Maximum residual topographies are between 7850 and 6700 m, whereas the minimum are between −3770 and −2420 m. Second, both models evidence that the cratonic areas in the northeast South America, including the whole northern transect analyzed in this work, are characterized by large positive residuals, suggesting an overcompensation state (dynamic uplift?). The Pampas to Patagonian pericratonic areas, in the southernmost South America, show contradictory results. While with Lithos 1.0 data (Fig. 1B) residual topography shows negative values, with Priestley and McKenzie (2013) data and Crust 1.0 the residuals are positive (Fig. 1C). It is important to note that topographic computation using only the crustal thicknesses from these models explains almost the entire continental South American elevations. It means that, except in some regions, the continent is close to crustal equilibrium. Then, the addition of a heavier lithospheric mantle produced by the Earth surface displaces under sea level. This is clearly shown in the cratonic areas, where lithospheric mantle can reach 150 km (or more). It is important to note that our results agree with Hu et al. (2018) model.

4.2 ­Local residual estimations 4.2.1 ­Paleolithosphere results and Cretaceous isostatic elevation

The lithospheric conditions during the Cretaceous in the central transect of the south-central Andes can be well constrained given the abundant information available from the alkaline magmatism that characterizes the region, and which shows some similarities with the Paraná-Etendeka Province (Lucassen et al., 2007). According to the geochemical data, they are OIB-type rocks that originated from a low melting degree in the garnet stability field of the mantle (Lagorio, 2008). Both depleted and enriched mantle compositions were recorded (Lucassen et al., 2007). Thermobarometric calculations performed with samples from the Sierras de Córdoba (Kay and Ramos, 1996; Lucassen et al., 2007; Lagorio, 2008) allowed us to calculate a lithospheric thickness of 76 ± 6 km and high-temperature conditions compatible with an origin from a mantle plume. We also obtained a similar value using the rare earth elements inverse modeling of Gibson and Geist (2010), which yields a lithosphere 71 ± 1 km thick using the same database. Previously, Lucassen et al. (2005) had calculated a lithosphere of approximately 70 km thick based on thermobarometric studies of peridotite xenoliths included within the Cretaceous magmas of NW Argentina, reinforcing our interpretations. The calculated thicknesses are also compatible with the deep melting conditions, described previously (Kay and Ramos, 1996; Lucassen et al., 2007; Lagorio, 2008). With these paleolithosphere thicknesses and following the methods described in Section 3.1 (residual topography) and using the density contrasts that result from the topographic balance analysis to reproduce the modern elevations in this region (i.e., density test to reproduce the observed modern topography, considering the modern lithosphere thickness is less than the dynamic contribution that results from backstripping far away from Andean loads. This yielded 3200 kg/m3 for the asthenosphere, 3280 kg/m3 for the lithospheric mantle, and 2800 kg/m3 for the crust), we obtained a CretaceousPaleogene paleoelevations between 1.4 and 1.5 km above sea level, which contrasts with contemporaneous marine records (Malumian and Nanez, 2011; Reinante et al., 2014; Marengo, 2015) and the ancient changes of dynamic topography in the Cretaceous (see red and blue curves in Fig. 6B).

52

Chapter 3  Measuring dynamic topography in South America

4.2.2 ­Subsidence analysis ­Northern-Central Andes and Amazon foreland transect

­Flexural. Our previous estimation (Dávila and Lithgow-Bertelloni, 2013) in the Peruvian foreland was based on reconstructions considering only the Eastern Cordillera and Subandean thrust belt loading, between 12° and 14°S (see Gotberg et al., 2010). However, this approach underestimated the loading effect of the western belts, like the Western Cordillera (known as the Marañon thrust belt: morphotectonics cf. Benavides-Cáceres, 1999). Recent shortening estimation of this area at ~10°S in the Peruvian Andes (see Scherrenberg et al., 2014, 2016) will allow polishing and evaluating our previous subsidence history interpretations. Given that both sections (i.e., Gotberg et al., 2010; Scherrenberg et al., 2014) are not physically continuous, we construct a composite broken load reconstruction with a minimum separation between them in order to enhance the influence of tectonic loads on the distal foreland (see Fig. 5A). It is important to note that we do not consider the influence of small intraforeland blocks, such as Shira-Othisi. However, like in the broken foreland of Argentina, these elements seem to have very little influence on the regional loading and subsidence (see Dávila et al., 2010). The balanced cross-sections in the Marañon thrust belt, across the Western Cordillera at these latitudes, account for 38% of Andean deformation across ~70 km, which are amounts rather similar (30%–40%) to previous the orogenic shortening estimated to the east, in the Subandean Peruvian Andes (cf. Mégard, 1987; Hermoza et al., 2005) across ~180 km. It is important to clarify that this updated reconstruction does not account for the Eastern Cordillera shortening. Now the loading segment is ~250 km and the maximum subsidence is between 4400 and 4900 m (depending on the Te used, 20, 30, or 40 km) in the foredeep close to the loading and the extension of the subsidence area (to the forebulge) is ~100 km (see Fig. 5A). Likewise in our previous works, we compare the flexural results with the reported stratigraphic thicknesses across foreland basin sections (e.g., Espurt et al., 2007; Latrubesse et al., 2010). While Espurt et al. (2007) documented 1000–1500 m of Cenozoic thicknesses close to the Cordillera, 400 km cratonward (eastward), Latrubesse et al. (2010) reported comparable 1100 m (Acre basin). This evidences a rather tabular geometry. Although the proposed loads can reproduce more than enough sedimentary thicknesses, the flexural model cannot explain the hundreds of meters of strata reported from 400 to >1000 km from the loads, from the Acres to Amazon basins. This mismatching agrees with our previous studies (Dávila and Lithgow-Bertelloni, 2013; Eakin et al., 2014), supporting that it is rather unlikely the Andean loads at these latitudes explain the subsidence in the distal pericratronic foreland of Peru and Brazil. ­Backstripping. In order to understand the subsidence in pericratonic areas of the Solimoes and Amazon basins (see Fig. 2), located out of influence of the Andes, we performed a backstripping analysis using the three of the westernmost boreholes extracted from Caputo and Soares (2016) and Heimdal et al. (2018) (see Fig. 6A). They record 2618 m of Paleozoic, Mesozoic, and Cenozoic sedimentary rocks. Results are shown in Fig. 6A and evidence a remarkable flat nonsedimentary driven subsidence since Upper Paleozoic. This subsidence is, however, nearly a hundred of meters from the beginning of the Cenozoic, evidencing that extra loading is needed.

­Southern-Central Andes—Precordillera-Pampas foreland transect

­Flexural. We focus our flexural studies in the Argentine Precordillera belt on the base of new structural reconstructions (Sánchez et al., 2019) that considers, in contrast to previous (e.g., Allmendinger and Judge, 2014), intermediate shortening stages between the middle Miocene to Present. To date, the associated basin (or Bermejo foreland) history was modeled using a total tectonic load, i.e., considering the whole accumulated shortening between Miocene to Present (from an undeformed to the present-day

4 ­  Results

53

shortened setting, see Cardozo and Jordan, 2001; Dávila et al., 2007; Fosdick et al., 2015). We compare our flexural subsidence results with the subsidence history derived from stratigraphic studies (Jordan et al., 2001) (Fig. 5C). Our results show (in support of our previous conclusions) a strong mismatching between flexural and subsidence curves since the Middle Miocene (Fig. 5B). None loading stage from 14 to 0 Ma correlates with the stratigraphic subsidence curves (Fig.  5C). Our results reaffirm the need for complementary loads, whether they are lithospheric or sublithospheric. Although we have not modeled the Cordilleran loads (Principal and Frontal Cordilleras regions) given that the shortening and tectonic loading would have occurred here before, the mismatching with coeval sedimentary sequences located further East (far away from the Cordilleran influence, in the Sierras Pampeanas) would reinforce the Dávila et al. (2007) hypothesis related to the influence of mantle forces in the foreland configuration (see also Dávila and Lithgow-Bertelloni, 2013). However, as discussed later, these results are in conflict with the dynamic uplifting proposals occurring after the slab flattening in the early late Miocene to this region. ­Backstripping. We modeled two different regions across this transect, in the foredeep area, where the maximum Cenozoic thicknesses were reported (e.g., Milana et al., 2003) and hundreds of kilometers eastward, in the Pampean pericratonic foreland, where tectonic loading is null (Fig. 6B). In the Bermejo foredeep, we use one borehole (Bermejo x-1) information, represented by ~5000 m of Cenozoic clastic strata. Stratigraphic unit information is summarized in Milana et al. (2003). The nonsedimentary-driven subsidence in this region is strongly steep, showing values of approximately 1700 m, which show a trend comparable to the flexural curves described in the Precordillera (Fig. 6B in dash line). Toward the Pampas, we modeled five boreholes (Ordoñez, Nogoya, Levalle, Camilo Aldao, and Las Toscas). Stratigraphy of this region is synthesized by Reinante et al. (2014). These boreholes show curves relatively steep during the Mesozoic, which become flatter. However, likewise in the Amazonas boreholes, the subsidence histories show small (but significant) slopes. This computation shows that ~100 m of subsidence is not related to tectonic and/or sedimentary loading. The pericratonic setting drives us to suggest subcrustal forces (see discussion below). This dynamic subsidence component in the Pampean area is the one used for the isostatic elevation calculations and then to adjust crustal and mantle densities to account for the modern altitudes (between 150 and 0 m on sea level). These are the densities used in paleoisostatic elevation studies (e.g., Cretaceous, see earlier).

­Southern Andes—Patagonian transects

­Flexural. Recent structural reconstructions in Patagonian Southern Andes (see Ghiglione et al., 2014) allow us to improve our previous comparisons between flexural models and basin records (Dávila and LithgowBertelloni, 2013). According to this work, although shortening is rather similar to previous estimations, load distribution is different (see Fig. 5D). Our new results show three likely alternatives, based on Te values. Maximum subsidence, close to loads, are between 1130 and 1075 m and the basinal region varies between 115 and 215 km. When these values are compared with basin geometries (Ramos, 1989; Ghiglione et al., 2010), residual spaces arise, suggesting additional subcrustal loads to account flexural bending. ­Backstripping. This study is based on 10 oil boreholes from the Golfo San Jorge and Austral basins, located hundreds of kilometers away from the easternmost tectonic load (Fig. 6C and D) The sedimentary record in both areas is over 2000 m deep, formed by Mesozoic and Cenozoic clastic units (see stratigraphic details in Ghiglione et al., 2010). Boreholes show backstripping curves relatively steep during the Mesozoic, in consistency with synrifting, which become flatter during the Cenozoic. However, likewise in the boreholes reported toward the north, subsidence show small but significant slopes.

54

Chapter 3  Measuring dynamic topography in South America

At least 200–400 m of subsidence is not related to sedimentary and/or thrust-tectonic loading. According to the setting, it should be associated with subcrustal forces.

5 ­Discussion, ongoing studies, and future perspectives Our preliminary residual topography computations as well as the revision of the most recent work on topographic analysis at large scales allow us reasserting and supporting our hypothesis on dynamic forces in the construction of the observed topography of South America (Dávila and LithgowBertelloni, 2013, 2015). The residual remaining elevations signal that results from an observed (or total) and model topography can be viewed as equivalent of the “observed or measured” dynamic components. Although this residual topography is likely the closest-fit proxy to measure dynamic subsidence or uplift, the solutions sometimes can be ambiguous and complex. While the models considering an average elevation (e.g., our Fig. 1A, see Flament et al., 2013) are too simplistic and disregard the influences of lithospheric changes in time (crustal or lithospheric mantle thickening or thinning), the alternative by accounting for the isostatic contribution to surface topography is tricky. An accurate knowledge of the density structure and thickness of both the crust and lithospheric mantle is required to calculate the isostatic topography. In fact, different results might account from different acknowledged models. Our Fig. 1B and C are clear examples. Then, a question arises: What is the best residual topography fitting to constrain dynamic topography? A solution is a crossed analysis using other potential proxies of residual topography. Here we propose that residual topographies can also result from more local studies, in complement to lithospheric residual computations, like 1D and 2D subsidence analyses, e.g., backstripping and/or flexural models. In South America, these studies have shown us that the sedimentary and tectonic loads alone can explain (totally or partly) the very thick stratigraphic sequences preserved in foredeeps, i.e., close to the Andes, but are very insufficient to generate accommodations far away (>200 km) within the distal foreland to cratonic-platform areas (see Dávila and Lithgow-Bertelloni, 2013, 2015; Eakin et al., 2014; and our Figs. 6 and 7). Approaches of paleoelevation studies also evidence that crustal thickening by deformation is not always the answer to generate the modern relief (e.g., Hoke et al., 2014; Rodríguez Tribaldos et al., 2017; Nóbile, 2013). These studies have evidenced local “residual topographies” that requires subcrustal to sublithospheric processes to reproduce the modern topography (Fig. 7). A classic example (no evaluated in this work) is the late Miocene to Present rising of the Altiplano-Puna plateau in the core of the Central Andes, related to lithospheric thinning (see Garzione et al., 2017). Rodríguez Tribaldos et al. (2017) performed a linear inverse modeling of a large inventory of rivers profiles across South America to calculate cumulative uplift history using Pritchard et al. (2009) and Roberts and White (2010) approaches. Their methodology is based on the simple assumption that drainage system is mainly controlled by erosion triggered by uplift, by choosing the adequate erosional parameters that set the pace of river profiles erosion, they can invert their uplift history. They constrain their results using independent geologic observations of regional uplift from three different areas along South America. Rodríguez Tribaldos et al. (2017) demonstrated that the bulk of South American topography grew constantly during the Cenozoic, particularly in the last 20 Ma. Although this uplift might correlate with the presence of low-density anomalies in the mantle, underlying the Brazilian cratons, some limitations (real densities for example) allow us to propose that part of the modern topography must be generated and supported dynamically. Our more local studies using inverse modeling of single river profiles (Nóbile, 2013; Ávila, 2015) show similar uplift, with older uplift pulses to the north (Eocene to Miocene) and more recent

FIG. 7 (A–D) Uplift analysis in the south-Central Andes using paleoelevation studies (cf. Hoke et al., 2014), and (E) low-temperature thermochronology (based on data compilation, apatite fission track to the left and apatite helium to the right). The figures on top show: (A) Crustal and (B) lithospheric mantle thicknesses, (C) isostatic elevations, and (D) residual topographies that derive from the previous models. Notice that the maximum elevation changes proposed by Hoke et al. do not match with the residuals. Sample location of Hoke et al. (2014) is shown as reference.

56

Chapter 3  Measuring dynamic topography in South America

uplift episodes toward the southeast along the Sierras de Córdoba (Miocene) and northern Patagonia, along the Colorado River (<2 Ma). Our different residual estimates along the northernmost transect in the cratonic and pericratonic areas, within the Amazon foreland (far away from the Peruvian Andean loads), suggest dynamic subsidence eastward from the leading edge of the flat slab at least from the Miocene. Both studies, backstripping (Fig. 6) and flexural approaches (Fig. 5), are consistent and agree with our recent dynamic topography models (e.g., Eakin et al., 2014, see also Fig. 4). Bishop et al. (2018) recently developed seismic velocity studies across the flat-slab segment of Peru and suggested that dynamic uplifting is not necessary to explain the elevations across this area, and particularly to reproduce the Fitzcarrald Arch elevations (as proposed by Eakin et al., 2014). All this information, collected at more local scales (i.e., subsidence analyses and seismic velocity studies), however, disagrees strongly with our regional residual topography calculations, which show a very strong dynamic upwarping component across this transect (Fig. 1) and covering almost the whole north part of South America or, what it is the same, a very strong dynamic uplift. This large-scale residual component is likely the clearest of South America, given that all lithospheric models (CRUST1.0 + Priestley and McKenzie, 2013, and LITHOS1.0 data) reproduce the same positive signals (compare Fig. 1B and C); even using very low-density contrasts. A question arises: How to compatibilize our results, i.e., creation of sedimentary accommodation spaces during uplifting? Our work suggests that a thick mantle root drives surface sinking below sea level and, therefore, upwarping dynamic forces (or unreal densities) are needed to reach the present-day elevations above sea level. Hu et al. (2018) recently proposed that Cretaceous delamination below cratonic areas in Brazil might have driven to the formation of a lighter lithospheric mantle. These apparent incoherencies and inherited mechanisms from the Gondwana breakup are likely telling us about the processes that produce basins in uplifting scenarios (Fig. 8). While lithospheric residuals need of dynamic uplift, to keep Earth’s surface above sea level, lateral changes of dynamic forces might be generating localized basinal areas and, preservation of Cenozoic sequences very far away from the Andes within “intracratonic” continental areas (Fig. 8A). The final resultant is an elevated but localized depocenter, within a “stable” area. We might also think of a lateral constant dynamic force to keep the surface on sea level but with a lithosphere formed by a thickened lithospheric mantle root underlying the basin areas (Fig. 8B). In the flat slab segment of Chile-Argentina, where the Juan Fernández oceanic ridge subduction and slab flattening occur, regional and local residual topography data show enough consistent with dynamic topography results (own as well as others compiled from literature, see GPlate Portal online compilation). Likewise in Peru, our backstripping and flexural studies (Figs.  6 and 7) evidence dynamic subsidence across the foreland areas. But, in contrast to the northern transect, our regional residual topographies in the south-Central Andes (middle transect in Fig. 1B and C) evidence equilibrium to negative values, in agreement with local residuals and our previous dynamic topography models (see Fig. 4). It is important to highlight that our subsidence curve derived from time-step flexural analysis (Fig. 5) evidence a strong mismatching with the tectonic subsidence history derived from previous stratigraphic studies in the foredeep and our backstripping derived from the Bermejo borehole. In fact, the nonsedimentary subsidence in the thickest sections exposed in the northern Precordillera (Guandacol to Vinchina towns, where Miocene-Present is >10 km thick) required unreal tectonic loads to reproduce these spaces (Dávila et al., 2005). All these results are more compatible with a continuous dynamic subsidence since the middle Miocene to Present. However, these data are not compatible with the elevation changes reported by Dávila and Lithgow-Bertelloni (2015) in the Sierras Pampeanas and

5 ­ Discussion, ongoing studies, and future perspectives

57

FIG. 8 Three conceptual models to explain observations and residual topography models: (A) Lateral variations of dynamic topographic forces in order to account (at the same time) for elevations on sea level and dynamic subsidence and (B) Constant dynamic forces to account for elevations on sea level, with thickened lithospheric mantle roots to explain localized sedimentary accommodation spaces. Note that the dashed lines are the isostatic state. (C) This diagram shows the opposite situation, likely occurring in Patagonia, i.e., a dynamic downward force (dynamic subsidence) to reduce the isostatic elevation respect to observations (DEM).

the paleoelevation approaches (using isotopic δ18O values of precipitation at the sampling sites and the modern elevations) in Cordilleran belt by Hoke et al. (2014), particularly since the arrival of flat subduction. Our previous hypothesis stated that slab flattening drives dynamic uplift (Dávila and LithgowBertelloni, 2015). Fig.  8 shows the (a) crustal and (b) lithospheric thicknesses along this transect between 30° and 35°S, (c) our modeled isostatic topography that results from these layers, depicting the location of samples of Neogene basin records studied isotopically along this transect to understand paleoelevation (arrows are not to scale) and (d) the residual topography estimated like in our methodology. The southernmost samples, collected out of the influence of flat subduction and where the crust today is thinner than to the north, evidence a clear change of elevation and surface uplift of 1.2–2 km between the end of deposition (during the late Miocene, around 10 Ma) and present day (see Hoke et al., 2014). To the north, in contrast, where flat subduction develops at least since the Early Miocene (see Kay and Mpodozis, 2002) and the crust is thicker today, the paleoelevation changes are negligible to null. According to Hoke et al. (2014), the elevation changes cannot be linked to crustal shortening; therefore, the spatial pattern of surface uplift may be related to lower crustal flow. They d­ iscarded the influence of subduction-driven dynamic forces. However, it is important to note that Hoke et al. (2014) work considers an instantaneous modern picture in their analysis rather than changes since the earlymiddle Miocene to Present. If we do it, changes of dynamic topography might give a likely explanation to the reported changes of paleoelevation. We agree that the modern instantaneous dynamic topography on this flat-slab segment cannot reproduce this uplift (Sánchez et al., 2019) even less the negative values generated for other models (e.g., Müller et al., 2018). However, the changes of dynamic topography might have been significant (see Fig. 4C). But our previous models have not taken into account the southward slab flattening migration, in coincidence with the migration in the same sense of the Juan Fernández ridge subduction (Kay and Mpodozis, 2002), which might generate a similar trend in the change of dynamic topography. Taking into account the Hoke et al. (2014) considerations, we disagree

58

Chapter 3  Measuring dynamic topography in South America

with their connection with a likely lateral variation of exhumation. Our compilation of considerable amount of thermochronological ages (877 apatite fission track and 470 Uranium‑thorium‑helium dating, see Fig.  7E) does not show a clear along-strike anomaly in low-temperature cooling ages. An alternative explanation (or complementary) to the Hoke et al. (2014) proposal is to think the crust was much thinner previous to 10 Ma to the south or, alternatively, that the lithospheric mantle was thicker and reduced when slab shallowing moved southward (lithospheric refrigeration by slab flattening, see Collo et al., 2017). These changes might also account for the reported surface uplift. To the north, the topographic equilibrium state suggested by Hoke et al. (2014), i.e., no changes in 10 Myrs, is evidence that geodynamics and tectonics would have kept stable since then. However, as demonstrated here, with our local residual topography calculations, dynamic forces affected the Bermejo foreland basin formation (Figs. 6 and 7). This apparent contradiction, dynamic subsidence during uplift, reinforces our previous assumption stated here, along the northern-Central Andes transect, on the functioning of topographic driving mechanisms. It is likely that lateral changes of dynamic uplift drove more localized subsidence and basin formation. Future studies assimilating geological information will assist to arrive to more conclusive results. In Patagonia, at the Chile oceanic ridge subduction latitudes (Fig. 2), our preliminary results open new questions and challenges in order to understand the topographic evolution of the southernmost South America. While to date, most works have agreed (Guillaume et  al., 2009, 2013; Dávila and Lithgow-Bertelloni, 2013, 2015; Dávila et al., 2018) that dynamic subsidence worked together to tectonic loading to accommodate foreland sequences during the early-middle Miocene, dynamic uplift would have been associated with the rock uplift and exhumation of the foreland system, out of the influence of the Andean thrusts, generated by the slab window formation since Middle Miocene to Present (see Dávila et al., 2018). This hypothesis, in fact, might also explain the bypassing stage during the deposition of the “Rodados Patagónicos” (Ghiglione et al., 2016). Our regional and local residual topography analyses, however, show contradictory signals. Our 1D and 2D subsidence model results (Figs. 6 and 7), using backstripping and updated flexurals, suggest the necessity (likewise in previous models) of dynamic subsidence during the deposition of the youngest foreland successions, particularly toward the pericratonic foreland where tectonic loads are negligible. Furthermore, the lithospheric structures do not generate isostatic topographies to reproduce positive residual topographies, in agreement with the expected dynamic uplift since arrival and northward migration of the Chile oceanic ridge (Fig. 1). This new regional residual signal disagrees with our previous approaches in the Patagonian plateau, synthesized in Dávila et al. (2018) using the plate kinematic reconstructions of Breitsprecher and Thorkelson (2009) (see Fig. 4D–F). Residual topography models of Patagonia (see Fig. 1B and C, as well as Ávila and Dávila, 2017 and Hu et al., 2018) indicate that the plateau should have placed theoretically higher than today, and that dynamic dragging components should have developed during the arrival of the seismic ridge subduction in Patagonia, which agrees with backstripping models but contradicts (as stated earlier) with geomorphological and geodynamic approaches. Fig.  8C shows a likely alternative to explain the apparent contradiction between observations and our residual models. Our ongoing studies along this Patagonian transect have the objective to explain these apparently contrasting observations and models (Avila, PhD project). Our first step, as recently highlighted by Ávila and Dávila (2017), is to improve our knowledge on the Patagonian lithosphere. The lack of reliable lithospheric models prevents a reliable residual topographic analysis. The Patagonian lithosphere has been poorly studied and thicknesses are interpolated and extrapolated from thousands of kilometers away. A recent oil borehole bottom temperature compilation (Ávila and Dávila, 2018) in southern Patagonia will allow us estimating heat flows and a better lithospheric thicknesses.

6 ­Conclusions

59

A different strategy to analyze the regional uplift changes is to observe longitudinal river profiles (e.g., Pritchard et al., 2009), particularly on areas poorly affected by tectonics like those exposed in the pericratonic and cratonic areas of South America (see Fig. 2). These studies assume that the shape of a river was controlled by the history of uplift rate (Pritchard et al., 2009), and that a large inventory of longitudinal rivers profiles allows estimating regional uplift (e.g., Czarnota et al., 2014). Recent calculation in South America (Rodríguez Tribaldos et al., 2017) as well as our more local river analyses in Sierras Pampeanas and Patagonia (Nóbile, 2013; Ávila, 2015; Canelo, 2015) indicates that distal foreland belts, located out of the influence of tectonics, suffered surface uplift since 15–10 millions of years ago likely by dynamic forces. It is important to notice that Rodríguez Tribaldos et al. (2017) took into account all the knickpoints or zones, even those generated by tectonic features. If we only consider the nontectonic knickpoints and zones, the calculated dynamic uplift would be younger, in the last 5–2 Myrs (Ávila, 2015; Canelo, 2015). Independent of the uplift timing, these data agree with our regional residual topography models (Fig. 1B and C), but they do not match with subsidence analyses (backstripping and flexural, Figs. 6 and 7). Likewise we proposed in the pericratonic and cratonic areas of Peru and Brazil to the north above, we are likely observing two different scales and rates of subcrustal topographic-driven mechanisms. Our work reaffirm our previous conclusions on the importance of the location of the subduction leading edge across segments affected by oceanic ridge subduction to reproduce subsidence in areas located far away from tectonic loading, in pericratonic to cratonic areas. This is not only clear for Miocene example, like those described in the Amazonas and nearby regions, but also in the Argentine Pampas and Patagonia, but also for ancient settings, like the Cretaceous of the Pampas and Patagonia (see Gianni et al., 2018a,b), where dynamic subsidence was invoked to explain the Maastrichtian marine flooding as well as accommodations far away from the coeval orogen. In Patagonia, for example, this component allows explaining the apparently contradictory correlation between high sea levels and continental paleoenvironments between 120 and 100 Ma, and vice versa afterward, when marine records develop during a sea level fall (Gianni et al., 2018a,b). This interpretation also allows the coexistence of attenuated lithospheres proposed by Fosdick et al. (2014) to account for low effective elastic thickness values (but evidencing above sea-level elevations) and marine records (Biddle et al., 1986). In the Pampas, our paleolithospheric thickness computations suggest a clear above sea-level elevation during the Cretaceous of the regions located today in the modern Pampean plain. However, stratigraphic studies have reported marine sedimentation. This requires of a sublithospheric force to drag the surface of a light lithosphere below sea level. But the geodynamic setting during the Cretaceous of the south-Central Andes would not have been affected by slab flattening. Future studies will give more insights to account for the marine incursions of the core of South America.

6 ­Conclusions Our regional and local studies along the three segments affected by the subduction of oceanic ridges (aseismic and seismic) in the northern Central Andes, southern Central Andes, and Patagonian Southern Andes, confirm our hypothesis and previous interpretations on dynamic topography in South America. Based on the present-day knowledge of the South American lithosphere, we have the necessity to appeal to sublithospheric forces to account for different topographic features as well as basin and uplift records. The comparison of paleoisostatic elevations, based on a lithosphere estimation using basalt geochemical approaches in the core of South America, with paleoelevation markers, like marine

60

Chapter 3  Measuring dynamic topography in South America

p­ aleogeographies (i.e., sea level), allowed us extrapolating the dynamic topography hypothesis back in time, particularly during the breakup of Gondwana in the Cretaceous of South America. However, a mass anomaly underlying these regions has not been reported in the literature (for example, flat slabs). In contrast to our previous approaches, we arrive at the conclusion that the lithospheric mantle is a key element to consider in future topographic approaches, from basin analysis to uplift computations. The location of the lithosphere-asthenosphere boundary reduces significantly the large differences of dynamic components as well as the residual topography derived from different approaches, sometimes irreal. This has strong influences for past settings. The lithospheric mantle has changed during the Andean evolution since the breakup of Gondwana and the beginning of westward motion of South America (Husson et al., 2012). This is clearly demonstrated from the lithospheric thicknesses estimated in the Sierras Pampeanas, central Argentina, which changed from 70 to 80 km thick (estimated in this work) for the Cretaceous to 110 km to present day. Another interesting point to highlight is the significance of using reference frameworks to analyze large-scale subsidence and uplift. Multidisciplinary approaches are increasingly in need. A marine bed alone cannot be used as a proxy to interpret an ancient lithospheric state. The Cretaceous example of Patagonia, Pampas, or the cratonic areas of Amazonas are didactic. Our contribution marks the great importance of performing geological and geophysical observations together with numerical modeling to interpret large-scale geodynamic processes. Numerical approaches alone, simple or very sophisticated, require of an observational proxy to fit the diverse and complex physical parameters (viscosities, densities, velocities, etc.). The ground wires are the Earth records. The diverse and contrasting results on dynamic topography and/or residual topography (see Flament et al., 2013 and this contribution in see Fig. 1) clearly exemplify this. In fact, as evidenced in the GPlate Portal (compilation of different analysis), there is not even a full agreement between models developed by the same research groups.

­Acknowledgments We appreciate funding from UNC (SECyT), FONCyT (PICT-2015-1092, PICT-E 2018-0392), and CONICET (PUE-2016-CICTERRA).

­References Allen, P.A., Allen, J.R., 2013. Basin Analysis: Principles and Application to Petroleum Play Assessment. John Wiley & Sons, p. 677. Allmendinger, R.W., Judge, P.A., 2014. The Argentine Precordillera: a foreland thrust belt proximal to the subducted plate. Geosphere 10 (6), 1203–1218. Aragón, E., Pinotti, L., Fernando, D., Castro, A., Rabbia, O., Coniglio, J., Aguilera, Y.E., 2013. The Farallon-Aluk ridge collision with South America: implications for the geochemical changes of slab window magmas from fore-to back-arc. Geosci. Front. 4 (4), 377–388. Ávila, 2015. Análisis de la topografía de la región Norte de Patagonia, Argentina (36°-40° LS), a lo largo del Río Colorado y sus afluentes principales. (Undergraduate dissertation). Unpublished, 100 Universidad Nacional de Córdoba. Ávila, P., Dávila, F.M., 2017. Topografía residual y alzamiento dinámico de Patagonia austral y su vinculación con la ventana astenosférica entre 45° y 50° S, Argentina. In: Actas XX Congreso Geológico Argentino, San Miguel de Tucuman, pp. 77–79.

­References

61

Ávila, P., Dávila, F.M., 2018. Heat flow and lithospheric thickness analysis on the Patagonian slab windows, southern South America. Tectonophysics 747–748, 99–107. https://doi.org/10.1016/J.TECTO.2018.10.006. Beaumont, C., 1981. Foreland basins. Geophys. J. Int. 65, 291–329. Benavides-Cáceres, V.E., 1999. Orogenic evolution of the Peruvian Andes: the Andean cycle. In: Skinner, B.J. (Ed.), Geology and Ore Deposits of the Central Andes. 7. Society of Economic Geologists Special Publication, pp. 61–107. Biddle, K.T., Uliana, M.A., Mitchum, R.M., Fitzgerald, M.G., Wright, R.C., 1986. The stratigraphic and structural evolution of the central and eastern Magallanes Basin, southern South America. Foreland Basins, 41–61. Bishop, B.T., Beck, S.L., Zandt, G., Wagner, L.S., Long, M.D., Tavera, H., 2018. Foreland uplift during flat subduction: insights from the Peruvian Andes and Fitzcarrald Arch. Tectonophysics 731, 73–84. Bjørnerud, M.G., Austrheim, H., Lund, M.G., 2002. Processes leading to eclogitization (densification) of subducted and tectonically buried crust. J. Geophys. Res. Solid Earth 107 (B10). Blisniuk, P.M., Stern, L.A., Chamberlain, C.P., Idleman, B., Zeitler, P.K., 2005. Climatic and ecologic changes during Miocene surface uplift in the southern Patagonian Andes. Earth Planet. Sci. Lett. 230 (1–2), 125–142. Braun, J., Robert, X., Simon-Labric, T., 2013. Eroding dynamic topography. Geophys. Res. Lett. 40 (8), 1494–1499. Breitsprecher, K., Thorkelson, D.J., 2009. Neogene kinematic history of Nazca–Antarctic–Phoenix slab windows beneath Patagonia and the Antarctic Peninsula. Tectonophysics 464 (1–4), 10–20. Campbell Jr., K.E., Heizler, M., Frailey, C.D., Romero-Pittman, L., Prothero, D.R., 2001. Upper Cenozoic chronostratigraphy of the southwestern Amazon Basin. Geology 29 (7), 595–598. Canelo, H., 2015. Patrones y tasa de alzamiento de las sierras de Córdoba y del sur de Santiago del Estero (29°-34° S, 66°-63° O). (Undergraduate dissertation). Unpublished, 136 pp Universidad Nacional de Córdoba. Caputo, M.V., Soares, E.A.A., 2016. Eustatic and tectonic change effects in the reversion of the transcontinental Amazon River drainage system. Braz. J. Geol. 46 (2), 301–328. Cardozo, N., Jordan, T., 2001. Causes of spatially variable tectonic subsidence in the Miocene Bermejo Foreland Basin, Argentina. Basin Res. 13 (3), 335–357. Charrier, R., Ramos, V.A., Tapia, F., Sagripanti, L., 2015. Tectono-stratigraphic evolution of the Andean Orogen between 31 and 37°S (Chile and Western Argentina). Geol. Soc. Lond. Spec. Publ. 399, 13–61. Ciccioli, P.L., Limarino, C.O., Marenssi, S.A., 2005. Nuevas edades radimétricas para la Formación Toro Negro en la Sierra de los Colorados, Sierras Pampeanas Noroccidentales, provincia de La Rioja. Rev. Asoc. Geol. Argent. 60 (1), 251–254. Collo, G., Dávila, F.M., Teixeira, W., Nóbile, J.C., Sant′ Anna, L. G., Carter, A., 2017. Isotopic and thermochronologic evidence of extremely cold lithosphere associated with a slab flattening in the Central Andes of Argentina. Basin Res. 29 (1), 16–40. Cordani, U.G., Milani, E.J.A., Thomaz-Filho, D.A., Campos, 2000. Tectonic Evolution of South America. 854 pp. Rio de Janeiro. Cuitiño, J.I., Fernicola, J.C., Kohn, M.J., Trayler, R., Naipauer, M., Bargo, M.S., Vizcaíno, S.F., 2016. U-Pb geochronology of the Santa Cruz Formation (early Miocene) at the Río Bote and Río Santa Cruz (southernmost Patagonia, Argentina): implications for the correlation of fossil vertebrate localities. J. S. Am. Earth Sci. 70, 198–210. Czarnota, K., Roberts, G.G., White, N.J., Fishwick, S., 2014. Spatial and temporal patterns of Australian dynamic topography from river profile modeling. J. Geophys. Res. Solid Earth 119 (2), 1384–1424. Dahlen, F.A., 1990. Critical taper model of fold-and-thrust belts and accretionary wedges. Annu. Rev. Earth Planet. Sci. 18 (1), 55–99. Dávila, F.M., Astini, R.A., 2003. Early middle Miocene broken foreland development in the southern Central Andes: evidence for extension prior to regional shortening. Basin Res. 15, 379–396. Dávila, F.M., Astini, R.A., 2007. Cenozoic provenance history of synorogenic conglomerates in western Argentina (Famatina belt): implications for Central Andean foreland development. GSA Bull. 119 (5–6), 609–622. Dávila, F.M., Carter, A., 2013. Exhumation history of the Andean broken foreland revisited. Geology 41 (4), 443–446. Dávila, F.M., Lithgow-Bertelloni, C., 2013. Dynamic topography in South America. J. S. Am. Earth Sci. 43, 127–144. Dávila, F.M., Lithgow-Bertelloni, C., 2015. Dynamic uplift during slab flattening. Earth Planet. Sci. Lett. 425, 34–43.

62

Chapter 3  Measuring dynamic topography in South America

Dávila, F.M., Astini, R.A., Jordan, T.E., 2005. Cargas subcorticales en el antepaís andino y la planicie pampeana: evidencias estratigráficas, topográficas y geofísicas. Rev. Asoc. Geol. Argent. 60 (4), 775–786. Dávila, F.M., Astini, R.A., Jordan, T.E., Gehrels, G., Ezpeleta, M., 2007. Miocene forebulge development previous to broken foreland partitioning in the southern Central Andes, west-Central Argentina. Tectonics 26 (5), TC5016. https://doi.org/10.1029/2007TC002118. Dávila, F.M., Lithgow-Bertelloni, C., Giménez, M., 2010. Tectonic and dynamic controls on the topography and subsidence of the Argentine Pampas: the role of the flat slab. Earth Planet. Sci. Lett. 295 (1–2), 187–194. Dávila, F.M., Lithgow-Bertelloni, C., Martina, F., Ávila, P., Nóbile, J., Collo, G., Sánchez, F., 2018. Mantle influence on Andean and pre-Andean topography. In: The Evolution of the Chilean-Argentinean Andes. Springer, Cham, pp. 363–385. Davis, D., Suppe, J., Dahlen, F.A., 1983. Mechanics of fold-and-thrust belts and accretionary wedges. J. Geophys. Res. Solid Earth 88 (B2), 1153–1172. Eakin, C.M., Long, M.D., 2013. Complex anisotropy beneath the Peruvian flat slab from frequency-dependent, multiple-phase shear wave splitting analysis. J. Geophys. Res. Solid Earth 118 (9), 4794–4813. Eakin, C.M., Lithgow-Bertelloni, C., Dávila, F.M., 2014. Influence of Peruvian flat-subduction dynamics on the evolution of western Amazonia. Earth Planet. Sci. Lett. 404, 250–260. Espurt, N., Baby, P., Brusset, S., Roddaz, M., Hermoza, W., Regard, V., Bolaños, R., 2007. How does the Nazca Ridge subduction influence the modern Amazonian foreland basin? Geology 35 (6), 515–518. Ezpeleta, M., Dávila, F.M., Astini, R.A., 2006. Estratigrafía y paleoambientes de la Formación Los Llanos (La Rioja): una secuencia condensada miocena en el antepaís fragmentado andino central. Rev. Asoc. Geol. Argent. 61 (2), 171–186. Figueiredo, J.J.J.P., Hoorn, C., Van der Ven, P., Soares, E., 2009. Late Miocene onset of the Amazon River and the Amazon deep-sea fan: evidence from the Foz do Amazonas Basin. Geology 37 (7), 619–622. Flament, N., Gurnis, M., Müller, R.D., 2013. A review of observations and models of dynamic topography. Lithosphere 5 (2), 189–210. Flament, N., Gurnis, M., Müller, R.D., Bower, D.J., Husson, L., 2015. Influence of subduction history on South American topography. Earth Planet. Sci. Lett. 430, 9–18. Fosdick, J.C., Romans, B.W., Fildani, A., Bernhardt, A., Calderón, M., Graham, S.A., 2011. Kinematic evolution of the Patagonian retroarc fold-and-thrust belt and Magallanes foreland basin, Chile and Argentina, 51° 30′ S. Bulletin 123 (9–10), 1679–1698. Fosdick, J.C., Graham, S.A., Hilley, G.E., 2014. Influence of attenuated lithosphere and sediment loading on flexure of the deep-water Magallanes retroarc foreland basin, Southern Andes. Tectonics 33 (12), 2505–2525. Fosdick, J.C., Carrapa, B., Ortíz, G., 2015. Faulting and erosion in the Argentine Precordillera during changes in subduction regime: reconciling bedrock cooling and detrital records. Earth Planet. Sci. Lett. 432, 73–83. Fromm, R., Zandt, G., Beck, S.L., 2004. Crustal thickness beneath the Andes and sierras Pampeanas at 30 S inferred from Pn apparent phase velocities. Geophys. Res. Lett. 31 (6). Gansser, A., 1973. Facts and theories on the Andes: twenty-sixth William Smith Lecture. J. Geol. Soc. 129 (2), 93–131. Garzione, C.N., McQuarrie, N., Perez, N.D., Ehlers, T.A., Beck, S.L., Kar, N., Lease, R.O., 2017. Tectonic evolution of the central Andean plateau and implications for the growth of plateaus. Annu. Rev. Earth Planet. Sci. 45, 529–559. Georgieva, V., Melnick, D., Schildgen, T.F., Ehlers, T.A., Lagabrielle, Y., Enkelmann, E., Strecker, M.R., 2016. Tectonic control on rock uplift, exhumation, and topography above an oceanic ridge collision: Southern Patagonian Andes (47° S), Chile. Tectonics 35 (6), 1317–1341. Ghiglione, M.C., Ramos, V.A., Cristallini, E.O., 2010. Fueguian Andes foreland fold and thrust belt: structure and growth strata. Andean Geol. 29 (1), 17–41. Ghiglione, M.C., Likerman, J., Barberón, V., Beatriz Giambiagi, L., Aguirre-Urreta, B., Suarez, F., 2014. Geodynamic context for the deposition of coarse-grained deep-water axial channel systems in the Patagonian Andes. Basin Res. 26 (6), 726–745.

­References

63

Ghiglione, M.C., Ramos, V.A., Cuitiño, J., Barberón, V., 2016. Growth of the southern Patagonian Andes (46–53 S) and their relation to subduction processes. In: Growth of the Southern Andes. Springer, Cham, pp. 201–240. Gianni, G.M., Echaurren, A., Fennell, L., Navarrete, C.R., Quezada, P., Tobal, J., Folguera, A., 2018a. Cretaceous orogeny and marine transgression in the southern central and northern Patagonian Andes: aftermath of a largescale flat-subduction event? In: The Evolution of the Chilean-Argentinean Andes. Springer, Cham, pp. 291–328. Gianni, G., Dávila, F.M., Echaurren, A., Fennell, L., Tobal, J., Navarrete, C., Quezada Pozo, P., Folguera, A., Gimenez, M.E., 2018b. A geodynamic model linking cretaceous orogeny, arc migration, foreland dynamic subsidence and marine ingression in southern South America. Earth Sci. Rev. 185, 437–462. Gibson, S.A., Geist, D., 2010. Geochemical and geophysical estimates of lithospheric thickness variation beneath Galápagos. Earth Planet. Sci. Lett. 300 (3), 275–286. Giménez, M.E., Dávila, F., Astini, R., Martínez, P., 2011. Interpretación gravimétrica y estructura cortical en la cuenca de General Levalle, Provincia de Córdoba, Argentina. Rev. Mex. Cienc. Geol. 28 (1), 105–117. Gorring, M.L., Kay, S.M., 2001. Mantle processes and sources of Neogene slab window magmas from southern Patagonia, Argentina. J. Petrol. 42 (6), 1067–1094. Gorring, M.L., Kay, S.M., Zeitler, P.K., Ramos, V.A., Rubiolo, D., Fernandez, M.I., Panza, J.L., 1997. Neogene Patagonian plateau lavas: continental magmas associated with ridge collision at the Chile triple junction. Tectonics 16 (1), 1–17. Gotberg, N., McQuarrie, N., Caillaux, V.C., 2010. Comparison of crustal thickening budget and shortening estimates in southern Peru (12–14 S): implications for mass balance and rotations in the “Bolivian orocline”. Bulletin 122 (5–6), 727–742. Guillaume, B., Martinod, J., Husson, L., Roddaz, M., Riquelme, R., 2009. Neogene uplift of central eastern Patagonia: dynamic response to active spreading ridge subduction? Tectonics 28 (2). Guillaume, B., Gautheron, C., Simon-Labric, T., Martinod, J., Roddaz, M., Douville, E., 2013. Dynamic topography control on Patagonian relief evolution as inferred from low temperature thermochronology. Earth Planet. Sci. Lett. 364, 157–167. Gurnis, M., Mitrovica, J.X., Ritsema, J., van Heijst, H.J., 2000. Constraining mantle density structure using geological evidence of surface uplift rates: the case of the African superplume. Geochem. Geophys. Geosyst. 1 (7). Hager, B.H., O'Connell, R.J., 1981. A simple global model of plate dynamics and mantle convection. J. Geophys. Res. Solid Earth 86 (B6), 4843–4867. Heimdal, T.H., Svensen, H.H., Ramezani, J., Iyer, K., Pereira, E., Rodrigues, R., Callegaro, S., 2018. Large-scale sill emplacement in Brazil as a trigger for the end-Triassic crisis. Sci. Rep. 8 (1), 141. Heine, C., Zoethout, J., Müller, R.D., 2013. Kinematics of the South Atlantic Rift. . arXiv preprint arXiv:1301.2096. Heintz, M., Debayle, E., Vauchez, A., 2005. Upper mantle structure of the South American continent and neighboring oceans from surface wave tomography. Tectonophysics 406 (1–2), 115–139. Hermoza, W., Brusset, S., Baby, P., Gil, W., Roddaz, M., Guerrero, N., Bolaños, R., 2005. The Huallaga foreland basin evolution: thrust propagation in a deltaic environment, northern Peruvian Andes. J. S. Am. Earth Sci. 19 (1), 21–34. Herzberg, C., Asimow, P.D., 2015. PRIMELT3 MEGA. XLSM software for primary magma calculation: peridotite primary magma MgO contents from the liquidus to the solidus. Geochem. Geophys. Geosyst. 16 (2), 563–578. Hoke, G.D., Giambiagi, L.B., Garzione, C.N., Mahoney, J.B., Strecker, M.R., 2014. Neogene paleoelevation of intermontane basins in a narrow, compressional mountain range, southern Central Andes of Argentina. Earth Planet. Sci. Lett. 406, 153–164. Hoorn, C., Wesselingh, F.P., Ter Steege, H., Bermudez, M.A., Mora, A., Sevink, J., Jaramillo, C., 2010. Amazonia through time: Andean uplift, climate change, landscape evolution, and biodiversity. Science 330 (6006), 927–931. Horton, B.K., 2018. Tectonic regimes of the central and southern Andes: responses to variations in plate coupling during subduction. Tectonics 37, 402–429. https://doi.org/10.1002/2017TC004624. Hu, J., Liu, L., Faccenda, M., Zhou, Q., Fischer, K.M., Marshak, S., Lundstrom, C., 2018. Modification of the Western Gondwana craton by plume–lithosphere interaction. Nat. Geosci. 11 (3), 203.

64

Chapter 3  Measuring dynamic topography in South America

Husson, L., Conrad, C.P., Faccenna, C., 2012. Plate motions, Andean orogeny, and volcanism above the South Atlantic convection cell. Earth Planet. Sci. Lett. 317, 126–135. Jordan, T.E., 1981. Thrust loads and foreland basin evolution, Cretaceous, western United States. AAPG Bull. 65 (12), 2506–2520. Jordan, T.E., Allmendinger, R.W., 1986. The Sierras Pampeanas of Argentina; a modern analogue of Rocky Mountain foreland deformation. Am. J. Sci. 286 (10), 737–764. Jordan, T.E., Allmendinger, R.W., Damanti, J.F., Drake, R.E., 1993. Chronology of motion in a complete thrust belt: the Precordillera, 30-31 S, Andes Mountains. J. Geol. 101 (2), 135–156. Jordan, T.E., Schlunegger, F., Cardozo, N., 2001. Unsteady and spatially variable evolution of the Neogene Andean Bermejo foreland basin. Argentina. J. S. Am. Earth Sci. 14 (7), 775–798. Kay, S.M., Mpodozis, C., 2002. Magmatism as a probe to the Neogene shallowing of the Nazca plate beneath the modern Chilean flat-slab. J. S. Am. Earth Sci. 15 (1), 39–57. Kay, S.M., Ramos, V.A., 1996. El magmatismo cretácico de las Sierras de Córdoba y sus implicancias tectónicas. In: Actas 13 Congreso Geológico Argentino y 3 Congreso de Exploración de Hidrocarburos, Buenos Aires. vol. 3. pp. 453–464. Kay, S.M., Ramos, V.A., Marquez, M., 1993. Evidence in Cerro Pampa volcanic rocks for slab-melting prior to ridge-trench collision in southern South America. J. Geol. 101 (6), 703–714. Kohn, M.J., 2007. Paleoaltimetry: Geochemical and Thermodynamic Approaches. vol. 66. Mineralogical Society of America. Lachenbruch, A.H., Morgan, P., 1990. Continental extension, magmatism and elevation; formal relations and rules of thumb. Tectonophysics 174 (1–2), 39–62. Lagabrielle, Y., Suárez, M., Rossello, E.A., Hérail, G., Martinod, J., Régnier, M., de la Cruz, R., 2004. Neogene to quaternary tectonic evolution of the Patagonian Andes at the latitude of the Chile triple junction. Tectonophysics 385 (1–4), 211–241. Lagorio, S.L., 2008. Early Cretaceous alkaline volcanism of the Sierra Chica de Córdoba (Argentina): mineralogy, geochemistry and petrogenesis. J. S. Am. Earth Sci. 26 (2), 152–171. Laske, G., Masters, G., Ma, Z., Pasyanos, M., 2013. Update on CRUST1. 0—A 1-degree global model of Earth’s crust. Geophys. Res. Abstr. 15, 2658. Latrubesse, E.M., Cozzuol, M., da Silva-Caminha, S.A., Rigsby, C.A., Absy, M.L., Jaramillo, C., 2010. The Late Miocene paleogeography of the Amazon Basin and the evolution of the Amazon River system. Earth Sci. Rev. 99 (3–4), 99–124. Lithgow-Bertelloni, C., Richards, M.A., 1998. The dynamics of Cenozoic and Mesozoic plate motions. Rev. Geophys. 36 (1), 27–78. Liu, L., Gurnis, M., 2008. Simultaneous inversion of mantle properties and initial conditions using an adjoint of mantle convection. J. Geophys. Res. Solid Earth 113 (B8). Lucassen, F., Franz, G., Viramonte, J., Romer, R.L., Dulski, P., Lang, A., 2005. The late Cretaceous lithospheric mantle beneath the Central Andes: evidence from phase equilibria and composition of mantle xenoliths. Lithos 82 (3), 379–406. Lucassen, F., Franz, G., Romer, R.L., Schultz, F., Dulski, P., Wemmer, K., 2007. Pre-Cenozoic intra-plate magmatism along the Central Andes (17–34 S): composition of the mantle at an active margin. Lithos 99 (3), 312–338. Mackaman-Lofland, C., Horton, B.K., Fuentes, F., Constenius, K.N., and Stockli, D.F. n.d. Mesozoic to Cenozoic retroarc basin evolution during changes in tectonic regime, southern Central Andes (31-33°S): insights from zircon U-Pb geochronology: J. S. Am. Earth Sci. 89, 299–318, doi:https://doi.org/10.1016/j. jsames.2018.10.004. Malumian, N., Nanez, C., 2011. The Late Cretaceous–Cenozoic transgressions in Patagonia and the Fuegian Andes: foraminifera, palaeoecology, and palaeogeography. Biol. J. Linn. Soc. 103 (2), 269–288. Marengo, H.G., 2006. Micropaleontología y estratigrafía del Mioceno marino de la Argentina: Las Transgresiones de Laguna Paiva y del “Entrerriense-Paranense”. vol. 1. Universidad de Buenos Aires, p. 123.

­References

65

Marengo, H.G., 2015. Neogene Micropaleontology and Stratigraphy of Argentina: The Chaco-Paranense Basin and the Península de Valdés. 217 p, Springer. Martina, F., Canelo, H., Dávila, F.M., de Hollanda, M.H.M., Teixeira, W., 2017. Mississippian lamprophyre dikes in western Sierras Pampeanas, Argentina: structural evidences of transtensional tectonics along the SW margin of Gondwana. In: XX Congreso Geológico Argentino, Tucumán, Argentina. Mathalone, J.M.P., Montoya, M., 1995. Petroleum geology of the sub-Andean basins of Peru. In: Tankard, A.J., Suárez Soruco, R., Welsink, H.J. (Eds.), Petroleum Basins of South America. vol. 62. AAPG Mem, pp. 423–444. McClay, K., 2011. Introduction to thrust fault-related folding. In: McClay, K., Shaw, J.H., Suppe, J. (Eds.), Thrustfault Related Folding. vol. 94. AAPG Mem, pp. 1–19. Mégard, F., 1987. Cordilleran Andes and marginal Andes: a review of Andean geology north of the Arica elbow (18° S). In: Circum-Pacific Orogenic Belts and Evolution of the Pacific Ocean Basin, pp. 71–95. Milana, J.P., Bercowski, F., Jordan, T., 2003. Paleoambientes y magnetoestratigrafía del Neógeno de la Sierra de Mogna, y su relación con la Cuenca de Antepaís Andina. Rev. Asoc. Geol. Argent. 58 (3), 447–473. Mpodozis, C., Ramos, V.A., 1989. The Andes of Chile and Argentina. In: Ericksen, G.E., Cañas, M.T., Reinemund, J. (Eds.), Geology of the Andes and Its Relation to Hidrocarbon and Mineral Resources. Earth Sciences Seriesvol. 11. Circum-Pacific Council for Energy and Mineral Resources, pp. 59–90. Müller, R.D., Hassan, R., Gurnis, M., Flament, N., Williams, S.E., 2018. Dynamic topography of passive continental margins and their hinterlands since the Cretaceous. Gondwana Res. 53, 225–251. Nóbile, J.C., 2013. Tectónica y sedimentación pliocena-cuaternaria en las Sierras Pampeanas Septentrionales (27°29°l), Argentina. (Doctoral thesis). Unpublished Universidad Nacional de Córdoba, Córdoba, Argentina, p. 243. Parras, A., Griffin, M., Feldmann, R., Casadío, S., Schweitzer, C., Marenssi, S., 2008. Correlation of marine beds based on Sr-and Ar-date determinations and faunal affinities across the Paleogene/Neogene boundary in southern Patagonia, Argentina. J. S. Am. Earth Sci. 26 (2), 204–216. Pasyanos, M.E., Masters, T.G., Laske, G., Ma, Z., 2014. LITHO1. 0: an updated crust and lithospheric model of the Earth. J. Geophys. Res. Solid Earth 119 (3), 2153–2173. Pedoja, K., Husson, L., Regard, V., Cobbold, P.R., Ostanciaux, E., Johnson, M.E., Weill, P., 2011. Relative sealevel fall since the last interglacial stage: are coasts uplifting worldwide? Earth Sci. Rev. 108 (1–2), 1–15. Plank, T., Forsyth, D.W., 2016. Thermal structure and melting conditions in the mantle beneath the basin and range province from seismology and petrology. Geochem. Geophys. Geosyst. 17 (4), 1312–1338. Priestley, K., McKenzie, D., 2013. The relationship between shear wave velocity, temperature, attenuation and viscosity in the shallow part of the mantle. Earth Planet. Sci. Lett. 381, 78–91. Pritchard, D., Roberts, G.G., White, N.J., Richardson, C.N., 2009. Uplift histories from river profiles. Geophys. Res. Lett. 36 (24). Ramos, V.A., 1989. Andean foothills structures in northern Magallanes Basin, Argentina. AAPG Bull. 73 (7), 887–903. Ramos, V.A., 1999. Las provincias geológicas del territorio argentino. Geol. Argent. 29 (3), 41–96. Ramos, V.A., 2005. Seismic ridge subduction and topography: foreland deformation in the Patagonian Andes. Tectonophysics 399 (1–4), 73–86. Ramos, V.A., Kay, S.M., 1992. Southern Patagonian plateau basalts and deformation: backarc testimony of ridge collisions. Tectonophysics 205 (1–3), 261–282. Räsänen, M.E., Salo, J.S., Jungnert, H., Pittman, L.R., 1990. Evolution of the western Amazon lowland relief: impact of Andean foreland dynamics. Terra Nova 2 (4), 320–332. Regard, V., Lagnous, R., Espurt, N., Darrozes, J., Baby, P., Roddaz, M., Hermoza, W., 2009. Geomorphic evidence for recent uplift of the Fitzcarrald Arch (Peru): a response to the Nazca Ridge subduction. Geomorphology 107 (3–4), 107–117. Reinante, S., Olivieri, G., Salinas, A., Lovechio, J., Basile, Y., 2014. La cuenca Chacoparaná: estratigrafía y recursos de hidrocarborus. In: XIX Congreso Geológico Argentino, Córdoba, Argentina, pp. 895–912. Roberts, G.G., White, N., 2010. Estimating uplift rate histories from river profiles using African examples. J. Geophys. Res. Solid Earth 115 (B2).

66

Chapter 3  Measuring dynamic topography in South America

Rodríguez Tribaldos, V., White, N.J., Roberts, G.G., Hoggard, M.J., 2017. Spatial and temporal uplift history of South America from calibrated drainage analysis. Geochem. Geophys. Geosyst. 18 (6), 2321–2353. Rossetti, D.F., Netto, R.G., 2006. First evidence of marine influence in the Cretaceous of the Amazonas Basin, Brazil. Cretac. Res. 27 (4), 513–528. Rudge, J.F., Roberts, G.G., White, N.J., Richardson, C.N., 2015. Uplift histories of Africa and Australia from linear inverse modeling of drainage inventories. J. Geophys. Res. Earth Surf. 120 (5), 894–914. Ruskin, B.G., Dávila, F.M., Hoke, G.D., Jordan, T.E., Astini, R.A., Alonso, R., 2011. Stable isotope composition of middle Miocene carbonates of the frontal Cordillera and Sierras Pampeanas: did the Paranaense seaway flood western and Central Argentina? Palaeogeogr. Palaeoclimatol. Palaeoecol. 308 (3–4), 293–303. Russo, R.M., Gallego, A., Comte, D., Mocanu, V.I., Murdie, R.E., VanDecar, J.C., 2010. Source-side shear wave splitting and upper mantle flow in the Chile Ridge subduction region. Geology 38 (8), 707–710. Sánchez, N., et al., 2019. Constraining erosion rates in thrust belts: insights from kinematic modeling of the Argentine Precordillera, Jachal section. Tectonophysics. Schellart, W.P., 2017. Andean mountain building and magmatic arc migration driven by subduction-induced whole mantle flow. Nat. Commun. 8 (1), 2010. Scherrenberg, A.F., Holcombe, R.J., Rosenbaum, G., 2014. The persistence and role of basin structures on the 3D architecture of the Marañón fold-Thrust Belt, Peru. J. S. Am. Earth Sci. 51, 45–58. Scherrenberg, A.F., Kohn, B.P., Holcombe, R.J., Rosenbaum, G., 2016. Thermotectonic history of the Marañón Fold–Thrust Belt, Peru: insights into mineralisation in an evolving orogen. Tectonophysics 667, 16–36. Schmidt, C.J., Astini, R.A., Costa, C.H., Gardini, C.E., Kraemer, P.E., 1995. Cretaceous Rifting, Alluvial Fan Sedimentation, and Neogene Inversion, Southern Sierras Pampeanas, Argentina. Shephard, G.E., Müller, R.D., Liu, L., Gurnis, M., 2010. Miocene drainage reversal of the Amazon River driven by plate–mantle interaction. Nat. Geosci. 3 (12), 870. Somoza, R., 1998. Updated azca (Farallon)—South America relative motions during the last 40 My: implications for mountain building in the central Andean region. J. S. Am. Earth Sci. 11 (3), 211–215. Spasojević, S., Liu, L., Gurnis, M., Müller, R.D., 2008. The case for dynamic subsidence of the US east coast since the Eocene. Geophys. Res. Lett. 35 (8). Tassara, A., Swain, C., Hackney, R., Kirby, J., 2007. Elastic thickness structure of South America estimated using wavelets and satellite-derived gravity data. Earth Planet. Sci. Lett. 253 (1–2), 17–36. Turcotte, D., Schubert, G., 2014. Geodynamics. Cambridge University Press. Webster, R.E., Chebli, G.A., Fischer, J.F., 2004. General Levalle basin, Argentina: a frontier lower Cretaceous rift basin. AAPG Bull. 88 (5), 627–652. Yáñez, G.A., Ranero, C.R., Huene, R., Díaz, J., 2001. Magnetic anomaly interpretation across the southern central Andes (32–34 S): the role of the Juan Fernández Ridge in the late tertiary evolution of the margin. J. Geophys. Res. Solid Earth 106 (B4), 6325–6345. Zoback, M.L., Mooney, W.D., 2003. Lithospheric buoyancy and continental intraplate stresses. Int. Geol. Rev. 45 (2), 95–118.

­Further reading Garcia-Castellanos, D., 2002. Interplay between lithospheric flexure and river transport in foreland basins. Basin Res. 14 (2), 89–104. Hohertz, W.L., Carlson, R.L., 1998. An independent test of thermal subsidence and asthenosphere flow beneath the Argentine Basin. Earth Planet. Sci. Lett. 161 (1–4), 73–83. Winterbourne, J., Crosby, A., White, N., 2009. Depth, age and dynamic topography of oceanic lithosphere beneath heavily sedimented Atlantic margins. Earth Planet. Sci. Lett. 287 (1–2), 137–151.