U N I V E R S I D A D D E C O N C E P C I Ó N DEPARTAMENTO DE CIENCIAS DE LA TIERRA 10° CONGRESO GEOLÓGICO CHILENO 2003

SEARCHING FOR CAUSES OF ANDEAN SEGMENTATION (15°-47°S): INTERPRETATION OF ELASTIC THICKNESS AND ISOTOPIC DATA, PLUS A PRELIMINARY 3D DENSITY MODEL

TASSARA, A., GÖTZE, H-J., LUCASSEN, F.

Institut für Geologische Wissenschaften, Freie Universität Berlin, Malteserstrasse 74-100, D-12249 Berlin, Germany [email protected]

1. INTRODUCTION The Andean Cordillera is a continuous feature along the western margin of that has been constructed by a simple common process of ocean- collision. However, at a continental scale the show a remarkable along-strike segmentation of topography, volcanism, tectonics, geological history, subduction geometry and deep structure (e.g. Jordan et al., 1983; Mpodozis & Ramos, 1989; Cahill & Isacks, 1992; Kley et al., 1999; Tassara & Yañez, in press). The central (15°-33.5°S) and southern (33.5°-47°S) segments of the Andes constitute an excellent area to study the causes of this segmentation. In response to a convergence vector that does not exhibit spatial variations in azimuth and speed, the Central Andes developed a huge mountain belt that starkly contrasts with the narrow and small Southern Andes (fig. 1).

This fact, added to the persistency of the Andean segments throughout hundred of millions years (e.g. Mpodozis & Ramos, 1989), suggests that this segmentation is mainly controlled by inherited rheological variations along the continental plate.

Here we analyse this hypothesis using two different approaches. First, a quantitative interpretation of elastic thickness estimates in the framework of a brittle-elasto-ductile rheology is presented. This exercise, constrained by isotopic data, indicates that along-strike variations in crustal composition are the main factor controlling the rigidity differences between the Central and Southern Andes. Secondly, the results of a preliminary three-dimensional (3D) density model between 15°-47°S are presented. This model is focused on the identification of first-order along- strike variations of Andean crustal structure (geometry and composition).

2. DESCRIPTION OF THE ALONG-STRIKE SEGMENTATION OF THE ANDES (15°- 47°S) The position of oceanic ridges and the morphotectonic units forming the continental margin between 15°S and 47°S are depicted in figure 1. This figure also presents proposed boundaries and nomenclature for the segmentation along the Andean margin (after Tassara & Yañez, in press). The boundaries of the Central and Southern Andes correspond to the intersection of oceanic ridges with the Peru-Chile trench. Along the northern and southern boundaries of the Central Andes, the intermediate depth (100-200 km) dip angle of the subducted

Todas las contribuciones fueron proporcionados directamente por los autores y su contenido es de su exclusiva responsabilidad. slab changes abruptly from “flat” (<10°) to “normal” (25-30°) (e.g. Cahill & Isacks, 1992).

Fig. 1. Morphotectonic units of western South American between 15°S and 47°S and oceanic ridges of the . Boundaries and nomenclature for the Andean segmentation after Tassara & Yañez (in press). Abbreviations in bold emphasize morphotectonic units commented in the text: ssa Sierras Subandinas, ap Altiplano, wc Western Cordillera, pn Puna, sbs Santa Barbara System, cc Coastal Cordillera, fc Frontal Cordillera, sp Sierras Panpeanas, pc Principal Cordillera, pgc Patagonian Cordillera. Thin lines denote 50 km contour lines of the subducted slab (Cahill & Isacks, 1992). Triangles are active volcanoes of the central (grey, CVZ) and southern (white, SVZ) volcanic zones.

Both the Central and Southern Andes can be subdivided into second-order segments. We describe here some aspects of this second-order segmentation to be discussed in the next sections.

The width of the orogen decreases along the Central Andes from more than 600 km at the Altiplano Plateau to less than 200 km at the Frontal Cordillera segment. This narrowing correlates with a decrease in Neogene crustal shortening and a systematic change of the active foreland deformation styles: thin skinned at the Sierras Subandinas, thick skinned at the Santa Barbara System, and basement uplift at the (e.g. Kley et al., 1999). South of 24°S, the subducted slab shows a sub-horizontal zone below 100 km depth that reaches a maximum width of ~450 km at 33°-34°S (e.g. Cahill & Isacks, 1992). The absence of asthenospheric mantle below the Frontal Cordilleran lithosphere precludes the occurrence of a magmatic arc. Thus, the Central Volcanic Zone (CVZ) of the Andes is restricted to the Altiplano and Puna segments. The CVZ is composed of andesitic-dacitic calcalkaline stratovolcanoes located on top of the Western Cordillera and bi-modal volcanic products deposited on the Altiplano-Puna (e.g. Allmendinger et al., 1997). The magmas of the CVZ have assimilated a relatively high degree of felsic crust during their evolution (e.g. Kay et al., 1999).

South of 33.5°S, elevations higher than ~4000 m decrease gradually along the Principal Cordillera segment to less than 2000 m at the Patagonian Cordillera segment. This decrease also correlates with north-south variations in the Neogene tectonic style: thick-skinned foreland tectonics linked with trench-oblique structures at the Principal Cordillera segment, and trench- parallel strike-slip deformation along the Liquiñe-Ofqui Fault Zone (LOFZ) at the Patagonian Cordillera segment (e.g. Folguera et al., 2002). South of 34°S the subducted slab has a normal intermediate dip angle. This drives the mantle magmagenesis linked to the presence of the Southern Volcanic Zone (SVZ) of the Andes. The northern part of the SVZ, on top of the Principal Cordillera, is composed of andesitic calcalkaline stratovolcanoes with a crustal magmatic component decreasing toward the south (e.g. Hildret & Moorbarth, 1988). The southern SVZ builds the western-central part of the Patagonian Cordillera and is genetically related to the LOFZ (e.g. Lopez-Escobar et al., 1995). This volcanic chain is composed of andesitic and basaltic stratovolcanoes of calcalkaline to tholeitic affinities and several minor eruptive centres.

3. ISOTOPIC DATA, ELASTIC THICKNESS AND ANDEAN SEGMENTATION

3.1. ISOTOPIC CHARACTERIZATION OF THE ANDEAN CRUST Figure 2 summarises new unpublished and published Sr, Nd and Pb isotopic data. Isotopic compositions of typical magmatic, metamorphic, and sedimentary rocks constrain the crustal make up in the Central Andes between 18° and 28°S and the Southern Andes between 36° and 40°S.

Fig. 2. Isotopic composition of Palaeozoic magmatic and metamorphic rocks (Palaeozoic crust), Meso-Cenozoic magmatic rocks and Neogene volcanic rocks from in the Central (18°-28°S, red symbols) and Southern (36°- 40°S, blue symbols) Andes. Data of pre-Neogene rocks are corrected for in-situ decay. Data sources: Patagonian Batholith, Pankhurst et al. (1999); Central Andes Jurassic – early Cretaceous, Lucassen et al. (2002); SVZ Cenozoic volcanic rocks, Hildreth and Moorbath (1988); Southern Andes pre-Neogene rocks, unpublished data of Lucassen (pers. comm.); All other data see Lucassen et al. (2001) and ref. therein.

From diagrams of figure 2, we observe the following points:

1. Isotopic composition of sampled Palaeozoic basement is very similar at both areas (fig. 3c-d). These Palaeozoic rocks are predominantly felsic in composition and represent the common crust of the Andean continental margin. 2. The felsic Palaeozoic crust in both areas has been compositionally modified by Meso- Cenozoic mafic-intermediate arc magmatism. The juvenile arc magmas came from a depleted mantle source (fig. 3a-b-e). 3. In the Central Andes, this process took place during the evolution of the nearly stationary Jurassic – early Cretaceous arc along the Coastal Cordillera. Later eastward migration of magmatism had only minor effect on the composition of the common Palaeozoic crust. Isotopic composition of the CVZ volcanic rocks shows a considerable contribution of the Palaeozoic crust (fig. 3b-e). 4. In the southern studied area, a continuous addition of juvenile magmas since ca. 200 My has occurred in a nearly stationary magmatic arc located in the Patagonian Cordillera. The new Meso-Cenozoic mafic-intermediate crust (Northpatagonian batholith) has largely an isotopic composition of depleted mantle (fig. 3a-b-e). Volcanic rocks of the SVZ are isotopically indistinguishable from the extended Meso-Cenozoic intrusions (fig. 3a). 5. In summary, along the axis of the present Andean magmatic arc, the bulk crustal composition revealed by isotopic data, changes from dominant felsic in the Central Andes (Palaeozoic crust) to mafic-intermediate in the Southern Andes (Northpatagonian batholith).

3.2. ELASTIC THICKNESS ESTIMATES The elastic thickness Te controls the flexural rigidity of the lithosphere. Thus, Te is a relevant parameter describing the long-term mechanical behaviour of the lithosphere (e.g. Burov & Diament, 1995). Tassara (1997; in press) and Tassara & Yañez (1996; in press, hereafter T&Y) present Te estimates for 15 profiles perpendicular to the trench axis along the Andean margin between 15° and 47°S. Forward modelling of the Bouguer anomaly was performed along each profile with the assumption that the continental lithosphere is a thin elastic plate of variable thickness (Te) deflected under topographic loads (full description in Tassara, 1997).

Fig. 3. After Tassara & Yañez (in press). a) Elastic thickness estimates across each modelled profile showing differences in Te throughout variations in colour and size. Moho depth Zm map was interpolated from values obtained after inversion of the modelled Bouguer anomaly for each profile. b) North-South profile A-A´ illustrates along- strike variations of elastic and crustal thickness. Figure 3a shows the final Te profiles of T&Y on a Moho depth (Zm) map interpolated from the inversion of the modelled Bouguer anomaly. The elastic thickness is low along the orogen axis (< 45 km) and high in the rigid forearc and foreland flanks (maximum Te ~70 km). This variation is consistent with results presented by other authors, in particular with those obtained by Kösters (1999) who applied the 2D coherence method to the southern Altiplano-Puna segment (see also Hackney et al., this volume).

Figure 3b is a north-south projection of the profile A-A´ defined in fig. 3a. This N-S profile intersects each modelled profile and illustrates variations of elastic thickness and Moho depths along the main orogen of the Central and Southern Andes. Values of Te < 10 km and Zm > 60 km characterize the Central Andean orogen. Along the Principal Cordillera, the elastic thickness increases and the crustal thickness decreases toward values characteristic of the Patagonian Cordillera, i.e. Te > 35 km, Zm < 40 km.

3.2.1. RHEOLOGIC INTERPRETATION OF TE ESTIMATES In order to give a quantitative rheological interpretation of the along-strike elastic thickness variations depicted in fig. 3b, we have developed a simplified version of the brittle-elasto-ductile approach described by Burov & Diament (1995). Figure 4 gives a summary of this method.

Fig. 4. Yield strength (σy(Z)) envelopes computed for weak quartzite (Qz, pink) and strong diabase (Db, green) -1 -n 2n -1 crusts, underlain (below Zm = 35 km) by a dunite mantle (Dn, blue). Creep parameters H [kJ mol ], n, A [N m s ] and thermal conductivity k [Wm-1 °K-1] for each material, are displayed in the table. These envelopes predict the mechanical behaviour (brittle, elastic or ductile) of crustal and mantle materials before loading. After a compressive load, simulated by an external stress σext = 100 MPa, the elastic behaviour is restricted to beams (grey areas) of thickness TeC=Zd[Qz]-Zb and TeM=Zd[Dn]-Zm for a Qz-crust decoupled from the mantle, and Te=Zd[Dn]-Zb for a coupled Db-crust. Expressions for brittle-elastic boundary Zb, elastic-ductile boundary Zd[Qz, Db or Dn] and “decoupled” elastic thickness TeD, are given in text.

The yield strength σy of a material at a given depth Z is the maximum external stress supported elastically before permanent deformation. Yield strength envelopes have been computed using Byerlee´s law (brittle deformation) and dislocation creep law (ductile deformation) for two end- member crustal materials: a weak quartzite (Qz, pink in fig. 4) and a strong diabase (Db, green in fig. 4). Both crusts are underlain, below a Moho depth Zm = 35 km, by a common dunite mantle (Dn, blue in fig. 4). Creep parameters A, n and H for each lithology (table in fig. 4) were taken from Burov & Diament (1995). As a first simplification we assume that temperature variation with depth is defined by a linear geothermal gradient that depends only on thermal conductivity k (table in fig. 4) and surface heat flow density Q (60 mW/m2 in this example). The yield strength envelope predicts the mechanical behaviour of crustal and mantle materials before external loading. Our second approximation is to assume that external stresses are defined by a constant arbitrary compressive stress σext. After loading, the lithosphere behaves elastically only in the depth range where σy(Z) ≤ σext (grey areas in fig. 4). In this context, the brittle-elastic boundary Zb at upper crustal level is defined by:

Zb = σext/B (1), where B is the compressive brittle gradient (here 25 MPa/km). The elastic-ductile boundary Zd at the base of each material layer (Qz, Db or Dn) is defined by:

−1/ n −1 Hk   e   Z Qz, Db, Dn = lnσ &  (2), d [] ext    nRQ   A 

-15 -1 where R and e& are the gas constant and strain rate (10 s in this example). When crust and mantle are mechanically coupled, i.e. Zd[Qz or Db] > Zm (Db-crust in fig. 4), the elastic thickness is defined by

Te = Zd[Dn] – Zb (3).

Following Burov & Diament (1995), if a ductile lower crust decouples the elastic beams of crust (TeC) and mantle (TeM), i.e. Zd[Qz or Db] < Zm (Qz-crust fig. 4), then the “decoupled” elastic thickness is defined by

3 3 3 3 3 3 Te D = TeC +Te M = ()Z d []Qz ∨ Db − Zb + (Z d []Dn − Z m ) (4).

Assuming σext = 100 MPa, taking the Zm values from fig. 3b, using realistic values of Q and e& (fig. 5a, constrained by published data when available) and computing for end-member crusts given by Qz and Db, we have used equations (1) to (4) to calculate two theoretical elastic thickness values for each point of the N-S profile defined in fig. 3b. The resulting Te profiles for Qz-crust and Db-crust are compared with the estimates of T&Y in fig. 5b.

This comparison shows that only a weak Qz-crust can reproduce the low elastic thickness of the Central Andes, while a strong Db-crust is necessary to explain the high Te values observed in the Patagonian Cordillera. Along the Principal Cordillera, a gradual southward increase in Db-crustal component is interpreted.

Fig. 5. In a) heat flow density values Q and strain rates e& along profile A-A´ (defined in fig. 3) used for the computations of theoretical elastic thickness values that define the coloured profiles of b). A Qz-crust (pink line) reproduce the Te values estimated by T&Y along the Central Andes (black line, as in fig. 3b), while a Db-crust (green line) does for the Patagonian Cordillera segment. See text for details.

3.3. GEODYNAMIC IMPLICATIONS This quantitative interpretation of elastic thickness estimates shows that the main factor controlling rigidity differences between the Central Andes and the Patagonian Cordillera is the composition (lithological constitution) of the continental crust: weak, quartz-rich, felsic crust in the Central Andes versus strong, plagioclase-rich, mafic-intermediate crust in the Patagonian Cordillera. These inferred variations in crustal composition are fully consistent with the isotopic data described in section 3.1. Thus, the addition of new strong and mafic Meso-Cenozoic crust to the felsic Palaeozoic basement along the stationary Patagonian arc should be considered as a fundamental process of Andean evolution. This is because it causes a first-order change in crustal rheology that dominates the future geodynamic development of the orogen. Under the same vigorous thermomechanic regime exerted since the Late Oligocene (e.g. Somoza, 1998), a weak felsic crust can easily absorb a significant fraction of the convergence (~20%, Hindle et al., 2002) by compressive deformation and shortening to form the Central Andes. In contrast, the strong mafic crust of the Patagonian Cordillera preclude the initiation of compressive deformation and the crust breaks and/or reuses margin-parallel fault zones (i.e. LOFZ). The concentration of magmatism along the LOFZ produce a further weakening of the structure and facilitates continued absorption of the trench-parallel component of convergence along the fault zone. The trench-orthogonal convergence component is taken up by pure subduction and overriding of the continental plate.

Along the entire Central Andean orogen, spatially homogeneous low elastic thickness (Te < 10 km) occurs in the presence of a drastic decrease of heat flow in the Frontal Cordillera segment with respect to the Altiplano-Puna segments (e.g. Hamza & Muñoz, 1996). Taking into account that the crustal thickness and composition of the Frontal Cordillera is similar for the northern segments, this north-south cooling is compensated, in terms of the low Te values, by a very low strain rate (see fig. 5a). This conclusion is supported by independent geologically- and geodetically-derived deformation rates (ref. in Tassara, in press and Tassara & Yañez, in press). This low deformation rate correlates with the along-strike decrease both in the width of the orogen and in total Neogene crustal shortening, and also with the increased basement involvement in the foreland deformation. If the weak crust of the Frontal Cordillera is limited to the east by a stronger crust related with the suspected Cuyania (Ramos et al., 1998 and refs. therein), the geodynamic evolution of the Central Andes could be controlled by an inherited southward narrowing of the weak felsic crust. In this context, and assuming that compressive deformation began at ~18 Ma simultaneously along the Central Andes (e.g. Jordan et al., 1997), the cessation of orogenic processes, caused by the increasing potential energy retained in the orogen, should occur first in the Frontal Cordillera and then in the northern segments. This is because of the lower total volume of weak crust that can be deformed. Geological data indicate that the uplift of the Frontal Cordillera was completed by Late Miocene (e.g. Ramos et al., 2002). Subsequently, deformation rates decreased and the deformation migrated to the east into the Sierras Pampeanas. This process would be, at least partially, responsible for the flattening of the oceanic slab. This is because the decrease of the convergence fraction absorbed by compressive deformation across the Frontal Cordillera lithosphere, with respect to the efficient thin-skinned shortening of the Altiplano segment, is compensated by more effective westward overriding of the continent above the slab.

4. PRELIMINARY RESULTS OF A 3D DENSITY MODEL Work in progress (A.T. Ph. D. thesis) aims to construct an integrated three dimensional (3D) density model for the entire Andean margin between 15° and 47°S. This 3D density model is produced by forward modelling of the measured gravity field using the software IGMAS developed at the Freie Universität Berlin (e.g. Schmidt & Götze, 1999). We will present an integrated model that includes an existing 3D density model constructed by Kirchner (1997) between 20° and 28°S, the model presented at this congress by Tašárová et al. (this volume) between 36° and 42°S, and a new model linking between both of these areas (i.e. between 28° and 36°S). This integrated model is focused on identifying first-order changes in deep crustal and mantle-lithospheric density structure along the Andean margin. In particular, the model will be used to study how the proposed compositional variations between the felsic crust of the Central Andes and the mafic crust of the Patagonian Cordillera can be resolved at depth and which structures control these variations along and across the Frontal and Principal Cordillera segments.

First-order constraints used to define the initial geometry along the new central part of the model include the geological map of Chile (scale 1:1.000.000, SERNAGEOMIN, 2002), the geological map of Argentina (scale 1:2.500.000, SEGEMAR, 1998), the geological map of South America (scale 1:5.000.000, Schobbenhaus & Bellizzia, 2001), crustal geological cross-sections proposed by Ramos et al. (2002), seismic hypocenter locations reported by Pardo et al. (2002a) and Campos et al. (2002), and a seismic tomography model presented by Pardo et al. (2002b).

The initial model defines the continental crust with 14 bodies that represent regional- to continental-scale geological units (basins, magmatic arcs, metamorphic belts, pre-Mesozoic , etc). The initial bodies will be later subdivided to account for internal inhomogeneities and density variations with depth. The bodies are organized so that they maintain a geologically consistent geometry that emphasises along-strike interfingering between them. We will show here the results of the final forward gravity modelling. However, the model can be completed only after compilation and processing of a homogeneous gravity data base (still in progress).

ACKNOWLEDGEMENTS The comments, discussions and exhaustive corrections of English language done by Ron Hackney substantially improved the final version of this manuscript. The presentation of this work is possible after the financial support of the Collaborative Research Project SFB 267 of the Deutscheforschungsgemeinschaft.

REFERENCES Allmendinger, R; Jordan, T; Kay, S; Isacks, B. 1997. The evolution of the Altiplano-Puna plateau of the Central Andes. Annual Reviews on Earth and Planetary Sciences, Vol 25, p. 139 - 174. Burov, E; Diament, M. 1995. The effective elastic thickness (Te) of continental lithosphere: What does it really mean?. Journal of Geophysical Research, Vol 100, No B3, p. 3905 - 3927. Cahill, T.; Isacks, B. 1992. Seismicity and shape of the subducted Nazca plate. Journal of Geophysical Research, Vol 97, No B12, p. 17503 - 17529. Campos, J., Hatzfeld, D., Madariaga, R., Lopez, G., Kausel, E., Zollo, A., Iannacone, G., Fromm, R., Barrientos, S., Lyon-Caen, H. 2002. A seismological study of the 1835 seismic gap in south-central Chile. Phys. Earth and Plan. Int. 132, p. 177-195. Folguera, A.; Ramos, V.; Melnick, D. 2002. Partición de la deformación en la zona del arco volcánico de los Andes neuquinos (36°-39°S) en los últimos 30 millones de años. Revista Geológica de Chile, V 29, N2, p. 227 – 240. Hamza, V.; Muñoz, M. 1996. Heat flow map of South America. Geothermics 25, p. 599 – 646. Hildreth, W; Moorbart, S. 1988. Crustal contribution to arc magmatism in the Andes of central Chile. Contribution to Mineralogy and Petrology, Vol. 98, p. 455 - 489. Hackney, R., Götze, H-J., Kirby, J. This volume. The Andean gravity field: important considerations before interpretation. Hindle, D., Kley, J., Klosko, E., Stein, S., Dixon, T., Norambuena, E. 2002. Consistency of geologic and geodetic displacements during Andean orogenesis. Geophysical Research Letters. Vol29, No8, 10.1029/2001GL013757. Jordan, T.; Isacks, B.; Allmendinger, R.; Brewer, J.; Ramos, V. 1983. Andean tectonics related to geometry of the subducted Nazca plate. Geological Society of America Bulletin, Vol 94, p. 341 - 361. Jordan, T.; Reynolds, J.; Erikson, J. 1997. Variability in age of initial shortening and uplift in the Central Andes, 16° - 33°30´S. In Tectonic Uplift and Climatic Change (W. Ruddiman Ed.) Plenum Press, p. 42 – 61. Kay, S.; Mpodozis, C.; Coira, B. 1999. Neogene magmatism, tectonism and mineral deposits of the Central Andes (22° - 33°S Latitude). In Geology and ore deposits of the Central Andes, Skinner, B. Eds, Society of Economic Geology Special Publication 7, p. 27 - 59. Kirchner, A. 1997. 3D-dichtemodellierung zur Anpassung des Schwere- und Schwerepotentialfeldes der zentralen Anden. Ph. D. Thesis (unpublished), Frei Universität Berlin, 98 p. Kley, J.; Monaldi, C.; Salfity, J. 1999. Along-strike segmentation of the Andean foreland; causes and consequences. Tectonophysics, Vol 301, p. 75 - 94. Kösters, M. 1999. 3D-dichtemodellirung des Kontinentalrandes sowie quantitative Untersuchungen zur rigidität der Zentralen Andean (20° - 26°S). Ph. D. Thesis (unpublished), Frei Universität Berlin, 181 p. Lopez-Escobar, L.; Cembrano, J.; Moreno, H. 1995. Geochemestry and tectonics of the Chilean southern Andes basaltic Quaternary volcanism (37° - 46°S). Revista Geológica de Chile, Vol 22, No. 2, p. 219 - 233. Lucassen, F.; Becchio, R.; Harmon, R.; Kasemann, S.; Franz, G.; Trumbull, R.; Wilke, H.; Romer, R.; Dulski, F. 2001. Composition and density model of the continental crust at an active continental margin – the Central Andes between 21° and 27°S. Tectonophysics, 341, p. 195 – 223. Lucassen, F., Escayola, M., Franz, G., Romer, R.L., Koch., K., 2002. Isotopic composition of Late Mesozoic basic and ultrabasic rocks from the Andes (23-32°S) - implications for the Andean mantle. Contributions to Mineralogy and Petrology 143, 336-349. Pankhurst, R.J., Weaver, S.D., Hervé, F., Larrondo, P., 1999. Mesozoic–Cenozoic evolution of the North Patagonian Batholith in Aysén, southern Chile. Journal of the Geological Society 156, 673-694 Mpodozis, C.; Ramos, V. 1989. The Andes of Chile and Argentina. In Geology of the Andes and its relation to hidrocarbon and mineral resources (G. Ericksen, M. Cañas-Pinochet & J. Reinemund Eds.), Circum-Apcific Council for Energy and Mineral Resources Earth Sciences Series, Vol 11, p. 59 - 90. Pardo, M.; Comte, D.; Monfret, T. 2002a. Seismotectonic and stress distribution in the central Chile subduction zone. Jour. South Am. Earth Sci. V15, p. 11 – 22. Pardo, M., Monfret, T., Vera, E., Eisenberg, A., Gaffet, S., Lorca, E., Perez, A. 2002b. Flat-slab subduction zone in Central Chile-Argentina: seismotectonic and body-wave tomography from local data. V International Symposium on Andean Geodynamics, Toulousse France, Extended abstracts, p. 469 – 473. Ramos, V.; Dallmeyer, R.; Vujovich, G. 1998. Time constraint on the Early Paleozoic docking of the Precordillera, central Argentina. In The proto-Andean margin of (Pankhurst, R.; Rapela, C. Eds.), Geological Society of London, Spetial Publication No 142, p. 143 – 158. Ramos, V.; Cristallini, E.; Pérez, D. 2002. The Pampean flat-slab of the Central Andes. Jour. South Am. Earth Sci. V15, p. 59 – 78. Schmidt, S. and H.-J. Götze, 1999: Integration of Data Constraints and Potential Field Modelling - an Example from Southern Lower Saxony, Germany. Physics and Chemistry of the Earth, Part A, Vol. 24, No. 3, pp. 191-196. Schobbenhaus & Bellizzia, 2001. Geologic Map of South America, scale 1:5.000.000. CGMW-CPRM-DNPM- UNESCO, Brazilia. SEGEMAR, 1998. Mapa Geológico de la República Argentina. Escala 1:2.500.000. Servicio Geológico Minero Argentina, Buenos Aires. SERNAGEOMIN, 2002. Mapa Geológico de Chile. Escala 1:1.000.000. Servicio Nacional de Geología y Minería, Chile. Carta Geológica de Chile, Serie Geología Básica N° 75, Santiago. Somoza, R. 1998. Updated Nazca (Farallon) - South America relative motions during the last 40 My: implications for the mountain building in the central Andean . Journal of South American Earth Sciences, Vol 11, No 3, p. 211 - 215. Tašárová, Z., Götze, H-J., Wienicke, S. This volume. Gravity data analysis and forward modelling along the Chilean Margin at 36-42°S. Tassara, A. 1997. Segmentación andina desde el análisis flexural de la anomalía de Bouguer. Memoria de título y tesis de magister (Inédito), Universidad de Chile, Departamento de Geología, 140 p. Tassara, A. In press. Nazca-South American plates interaction and the formation of the Central Andean Plateau: Review of a 2D flexural analysis along the Andean margin. Tectonophysics. Tassara, A., Yañez, G. en prensa. Rigidez de la litosfera andina (12°-47°S) y su relación con la segmentación geotectónica del orógeno. Rev. Geol. de Chile.