The Laramide fold-and-thrust belt in southern Mexico is characterized by N-S–trending structures in its central and eastern part and by NW-SE–trending structures in its western part. Here, we investigate, experimentally, the possibility that the Laramide structures of southern Mexico may be the result of inversion of previously thinned lithosphere zones under oblique compression. A revision of the geology of this region shows that the presence of two extensional basins, representing relatively weak blocks within more rigid lithosphere, strongly controlled the subsequent deformation pattern. For modeling purposes, we divided the southern Mexico lithosphere into blocks with different strength profiles: (1) a stable craton; (2) a weak block composed of the Guerrero Morelos Platform; (3) a relatively strong block exposing the pre-Cretaceous Tejupilco schist and the Early Cretaceous Teloloapan volcanic arc (Tejupilco anticlinorium); and (4) a weak block represented by the Arcelia–Palmar Chico basin. A series of physical experiments simulating the mechanical response of an analogue lithosphere composed of five simplified strength profiles was constructed. The model lithosphere was thinned orthogonally and shortened obliquely. Shortening was accommodated mainly by reactivation of preexisting extensional structures. The resulting orogenic deformation in the models is not entirely sequential and foreland-progressive. Inversion tectonics of extensional basins is thus proposed as an explanation for the structural diversity observed in Late Cretaceous shortening of southwestern Mexico. The predictions of our lithospheric model may be tested when more geophysical information about the structure of the southern Mexico lithosphere becomes available.


Late Cretaceous Laramide shortening deformation in southern Mexico occurred primarily through large-scale folds and thrusts involving Early Cretaceous and older rocks. These structures strike N-S in an E-W section of more than 160 km between Ciudad Altamirano and Papalutla (Cerca et al., 2007; Martini et al., 2009) and NW-SE farther to the west, NW of Zihuatanejo (Martini, 2008) (Fig. 1). The available geological information from southern Mexico consistently suggests the occurrence of an episode of lithosphere thinning prior to Late Cretaceous shortening (e.g., Centeno-García et al., 2008; Martini et al., 2009, and references therein). At a regional scale, at least two zones with geological evidence of Early Cretaceous extension can be identified from east to west (Fig. 1): (1) the Guerrero Morelos Platform, which has a sedimentary record of continental margin conglomerates and minor volcanism during Hauterivian–Aptian (Zicapa Formation), followed by a thick carbonaceous platform sequence during Albian–Turonian (Morelos Formation; Cerca et al., 2007); and (2) the Arcelia–Palmar Chico deep basin, which is characterized by a flysch-like sequence during the Valanginian–Aptian and an alternating sequence of limestone and dominantly mid-ocean-ridge basalt (MORB)–type volcanism during the Aptian–Coniacian (Arcelia–Palmar Chico Group; Martini et al., 2009). These two basins are separated by a block exposing the pre-Cretaceous metamorphic basement of the Tejupilco schist (Fig. 1) and the major Early Cretaceous Teloloapan volcanic arc. The scheme emerging from these first-order observations suggests that, as in many orogenic zones, inversion of extensional basins is an important mechanism in controlling the style and geometry of Laramide shortening. Furthermore, it may explain the occurrence of N-S–trending structures in the central and eastern part of the southern Mexico Laramide belt and coast-parallel, NW-SE–trending structures in the western part of the belt (Fig. 1C). In this work, we experimentally investigate the possibility that the N-S–trending contractile structures may have formed as a result of positive inversion tectonics during a phase of oblique compression.


Recent work has provided a wealth of new information on the geology of southern Mexico (Cerca et al., 2007; Talavera-Mendoza et al., 2007; Solari et al., 2007; Centeno-García et al., 2008; Martini et al., 2009; Mortensen et al., 2008). Southern Mexico geology is characterized by a finite number of different terranes or subterranes that can be recognized by a detailed analysis of the surface geology (Campa and Coney, 1983; Sedlock et al., 1993; Tardy et al., 1994; Dickinson and Lawton, 2001; Keppie, 2004). We focus our work on an ∼220-km-long and ∼70-km-wide, E-W–trending zone, with complete exposures of Cretaceous lithology and structures, located to the west of the Paleozoic Acatlan complex (Figs. 2 and 3).

The aforementioned papers provide a detailed reconstruction of the geological evolution of the region, which is characterized by an Early Cretaceous extensional phase in the Arcelia–Palmar Chico Basin and in the Guerrero Morelos Platform, predating the Laramide shortening. Based on these observations, we propose a model in which inversion tectonics played an important role in controlling the differences observed in the strain localization and deformation style observed in the Mesozoic lithostratigraphy of these terranes. Recent papers support a model in which the Mesozoic volcano-sedimentary successions of the Guerrero terrane were deposited directly on the thinned continental margin of the North American plate (Cerca et al., 2007; Centeno-García et al., 2008; Martini et al., 2009), leaving no space for accretion of allochthonous terranes as the cause for Late Cretaceous shortening. Thus, analogue models of lithospheric scale were constructed to explore the influence of lateral mechanical variations during successive phases of extension and compression on the structural style that resulted in the upper crust. The relative strengths of blocks used in the model were inferred from the basis of geological and deformation history.

The reader is referred to Cerca et al. (2007) and Martini et al. (2009) for a more extensive review of the geology of southern Mexico. In the following section, we discuss only the relevant data for construction of the model. Schematic stratigraphic columns and the relation between blocks and tectonic events are presented in Figure 3.

Arcelia–Palmar Chico–Huetamo Basin

The Arcelia–Palmar Chico Cretaceous basin was built on extended continental lithosphere and consists of the western Huetamo and eastern Arcelia–Palmar Chico areas, the lateral continuity of which has been recently confirmed by Martini et al. (2009). Cretaceous sequences were clearly deposited above metasedimentary rocks of Triassic age in the Huetamo area (Centeno-García et al., 2003, 2008; Talavera-Mendoza et al., 2007; Martini et al., 2009). Precambrian and Paleozoic detrital zircon ages and metamorphic continental clasts in conglomerates in the lower part of the Cretaceous sequence suggest that the region is underlain by a continental basement below the Arcelia–Palmar Chico sequence (Elías-Herrera et al., 2000; Talavera-Mendoza et al., 2007; Martini et al., 2009). Moreover, two large-scale batholiths of continental affinity are found on both sides of the Arcelia–Palmar Chico basin: the Placeres del Oro intrusive (ca. 119 Ma; Martini et al., 2009) and the Tingambato batholith (ca. 130 Ma; Garza-González Vélez, 2007; Martini et al., 2009).

The Cretaceous history of subsidence of the Arcelia–Palmar Chico basin starts in the Late Jurassic or Early Cretaceous with apron sedimentary deposits and volcanic rocks (Guerrero-Suástegui, 1997) of poorly constrained age (Angao Formation). A period of continuous subsidence and marine transgression in the late Valanginian to Aptian is recorded by the San Lucas Formation (Pantoja-Alor and Gómez-Caballero, 2003; Omaña-Pulido et al., 2005; Talavera-Mendoza et al., 2007; Martini et al., 2009), and there is a continuous record of marine sedimentation from the Valanginian to Hauterivian in the lower part of the Arcelia–Palmar Chico Group (Salinas-Prieto et al., 2000; Elías-Herrera et al., 2000).

A major episode of mafic volcanism in the Aptian to the Cenomanian (105–82 Ma; Delgado-Argote et al., 1992; Elías-Herrera et al., 2000) suggests localization of extension in the basin. The transition from pelagic limestone intercalated with mafic lavas in the Arcelia–Palmar Chico area (Salinas-Prieto et al., 2000; Elías-Herrera et al., 2000; Elías-Herrera, 2004) to platform and reefal limestone of Aptian age in the Huetamo area (El Cajon Formation; Pantoja-Alor, 1990; Skelton and Pantoja-Alor, 1999; Omaña-Pulido and Pantoja-Alor, 1998; Martini et al., 2009) and to subaerial conditions represented by deltaic clastic sediments and biostromic limestone of the Barremian–Early Aptian Comburindio Formation (Alencaster and Pantoja-Alor, 1998; Pantoja-Alor and Gómez-Caballero, 2003) indicates progressively shallower facies to the west, and implies a scenario of high subsidence rates (Elías-Herrera, 2004). This condition was established and lasted until the Albian–Early Cenomanian, when a period of carbonate deposition at the flanks of the basin is represented by the Mal Paso limestone in the Huetamo area (Pantoja-Alor, 1959; Buitrón-Sánchez and Pantoja-Alor, 1998; Pantoja-Alor and Skelton, 2000; Filkorn, 2002) and the Amatepec limestone in the Arcelia–Palmar Chico area (Elías-Herrera et al., 2000; Cabral-Cano et al., 2000).

The thick sequence of red beds of the Cutzamala Formation (Altamira-Areyán, 2002) partially covers the Arcelia–Palmar Chico basin, indicating a shift from marine to continental depositional systems. The base of the Cutzamala sequence is exposed at the flanks of the Tzitzio anticline, where a Late Cretaceous age has been established on a paleontological (Benammi et al., 2005) and geochronological basis (84 Ma for an interbedded lava flow—Mariscal-Ramos et al., 2005; ca. 74 Ma for an andesitic clast—Martini et al., 2009). Continental sedimentation was partly concurrent with the shortening and continued after it into the early Paleocene (Martini et al., 2009). The youngest red beds deposits unaffected by shortening are cut by late Eocene dikes and intrusive rocks (Serrano-Durán, 2005).

Shortening of the Arcelia–Palmar Chico basin involved a pervasive deformation, observed at the regional scale, in N-S–oriented folding and thrusting. Eastward thrusting of the Arcelia–Palmar Chico Group against the Tejupilco anticlinorium postdates the Cenomanian volcanic activity (Elías-Herrera et al., 2000). The volcanic units of the Arcelia–Palmar Chico Group and the Amatepec limestone are thrusted against the Tejupilco schist along two low-angle, parallel, N-S–oriented structures. Shortening in the Huetamo area is observed as large fold structures (Martini et al., 2009). In the northern part, folding is concentrated in the Tzitzio anticline, a major asymmetric fold with an axial plane steeply dipping to the west developed in post-Campanian times. The structural style of the Tzitzio anticline (ample folding in Late Cretaceous red beds) differs significantly from the pervasive folds and thrusts affecting the Arcelia–Palmar Chico area (thrusting and stacking of the Early Cretaceous sequence). Because of its Paleocene age, it has been interpreted as an out-of-sequence structure or as the product of a second phase of deformation (Martini et al., 2009).


The Tejupilco anticlinorium is a basement high between the Arcelia–Palmar Chico basin and the Teloloapan arc that exposes the Tejupilco schist (Elías-Herrera et al., 2000; Elías-Herrera, 2004). The youngest ages obtained for detrital zircons from the metasedimentary part of the Tejupilco schist are Late Triassic (Martini et al., 2009) and represent further evidence of an older continental crust below the Jurassic to Cretaceous sequences. Rock fragments of the Tejupilco schist have been recognized in debris deposits at the lower part of the Arcelia–Palmar Chico basin (Elías-Herrera et al., 2000). According to Elías-Herrera (2004), the original listric fault that put the Tejupilco schist in contact with the Palmar Chico Group was later reactivated as a fault propagation fold during inversion of the basin.

The Teloloapan volcanic arc (Campa and Coney, 1983; Salinas-Prieto et al., 2000; Talavera-Mendoza and Guerrero-Suastegui, 2000) is located in the eastern and southern part of the Tejupilco anticlinorium and was apparently constructed above continental basement (Salinas-Prieto et al., 2000; Elías-Herrera et al., 2000; Elías-Herrera, 2004). The main volcanic activity was Hauterivian–Aptian (Talavera-Mendoza and Guerrero-Suastegui, 2000), and detrital zircon ages in the associated sedimentary successions peak at ca. 129 and ca. 124 Ma (Talavera-Mendoza et al., 2007). Volcanism progressively waned between the Aptian and latest Albian (Guerrero et al., 1991; Monod et al., 2000). The variations in depositional environment recorded by ammonites found in the Teloloapan limestone suggest a rapid increase in depth from east to west (Monod et al., 2000), consistent with a westward migration and localization of the extension at the Arcelia–Palmar Chico basin at this time. The sequence is covered by the Late Cretaceous flysch-like sequence of the Pachivia-Miahuatepec Formation (Fries, 1960; Salinas-Prieto et al., 2000; Hernandez-Romano et al., 1997). Late Cretaceous to Tertiary red bed sequences are thin or absent in the Tejupilco anticlinorium, and the Eocene to Oligocene volcanic units were emplaced directly over the Tejupilco schist (Morán-Zenteno et al., 2007). This indicates that the block remained a topographic high from the Late Cretaceous onward.

Shortening deformation of the Cretaceous units of the Teloloapan arc was accomplished by two phases with opposite vergence (Salinas-Prieto et al., 2000; Cabral-Cano et al., 2000). The major structures have N-S orientation and eastward vergence, such as the low-angle Teloloapan thrust system, which juxtaposes the Teloloapan arc against the Guerrero Morelos Platform. Since there are no major thrust structures between the Tejupilco schist and the Teloloapan arc, for modeling purposes, we assume that they behaved as a single quasi-rigid block during shortening. The age of shortening is poorly constrained in this area because of the lack of stratigraphic markers, but it must have been post-Albian (the age of the limestone). Late Maastrichtian to early Paleocene postkinematic intrusive bodies emplaced along the front of the Teloloapan thrust system and farther to the east in the Guerrero Morelos Platform provide a minimum age for the end of thrusting (González-Partida et al., 2003; Meza-Figueroa et al., 2003). The structures with westward vergence are less intense and have been associated with back thrusting in the late stages of shortening (Salinas-Prieto et al., 2000).

Guerrero Morelos Platform Basin

The Guerrero Morelos Platform basin is a N-S–oriented, ∼100-km-wide zone with widespread exposures of Cretaceous platform carbonates framed between the Teloloapan arc anticlinorium and the Acatlan complex (Fig. 2; Fries, 1960; Hernández-Romano et al., 1997; Cerca et al., 2004, 2007). A Precambrian and Paleozoic basement has been inferred by Vélez (1990) and Levresse et al. (2004) to lie beneath the Guerrero Morelos Platform basin. The Guerrero Morelos Platform basin subsidence started in the Early Cretaceous, when over 1000 m of continental red beds, minor lava flows (40Ar-39Ar age of 127 ± 2 Ma; Fitz-Díaz et al., 2002), shallow marine to tidal anhydrites, and some limestone banks (Aptian-Albian) were deposited close to the cratonic border to form the Zicapa Formation and the Huitzuco anhydrite (Fries, 1960; de Cserna et al., 1980; Cerca et al., 2007). We interpret these sequences as the evidence of a period of large-scale extension and minor volcanism in the backarc of the active Teloloapan arc. Rapid extension was replaced in the late Albian by a period of steady subsidence, in which platform and reefal limestone developed in a platform environment (Morelos Formation; Fries, 1960; Hernández-Romano et al., 1997). Low and constant subsidence rates concluded by the end of the Cenomanian with the deposition of the overlying Mezcala Formation (Hernández-Romano et al., 1997). Limestone facies of the Morelos Formation are progressively shallower toward the east (Hernández-Romano et al., 1997), coastal and reefal facies directly overlie the Acatlan complex or its Jurassic cover (de Cserna et al., 1980), and limestone is very thin or completely absent over the eastern Acatlan complex, suggesting that the stable plateau of southern Mexico was only partially covered by the Albian-Cenomanian sea (Cerca et al., 2004, 2007).

The beginning of inversion in this area is marked by the switch from carbonate to siliciclastic sedimentation, marked by the Mezcala Formation (Fries, 1960; Ontiveros-Tarango, 1973; Hernández-Romano et al., 1997; Lang and Frerichs, 1998; Cabral-Cano et al., 2000; Cerca et al., 2004, 2007; Nieto-Samaniego et al., 2006). The Mezcala Formation ranges from the Turonian in the central part of the Guerrero Morelos Platform (Hernández-Romano et al., 1997), to the Coniacian in the Atenango del Rio area (Lang and Frerichs, 1998), to the early Maastrichtian in the Texmalac area (Perrilliat et al., 2000; Fig. 2). The end or at least waning of shortening is marked by the inception of subaerial magmatic activity and sedimentation since the end of Maastrichtian (González-Partida et al., 2003; Meza-Figueroa et al., 2003; Levresse et al., 2004; Cerca et al., 2007).

The red bed succession of the Guerrero Morelos Platform is similar in lithology, depositional environment, and age to the Cutzamala deposits in the Huetamo–Palmar Chico area. However, in the Guerrero Morelos Platform basin, continental deposits are less widespread and essentially confined to small structural basins (Cerca et al., 2004, 2007). The whole continental sequence is affected by gentle folding that tends to decrease toward the top (Cerca et al., 2004, 2007).

The eastern and western boundaries of the Guerrero Morelos Platform basin are N-S– and NE-SW–trending deformation zones, respectively: to the west, there is the Teloloapan thrust system, a low-angle, N-S–trending, and eastward-verging structure that juxtaposes the Early Cretaceous Teloloapan arc sequence with the Guerrero Morelos Platform limestone (Campa-Uranga, 1978; Campa-Uranga and Ramírez, 1979; Salinas-Prieto et al., 2000). To the east, the Guerrero Morelos Platform is bounded by the NE-SW–trending Papalutla fault, which thrust the Acatlan complex over the Guerrero Morelos Platform with a northwest vergence (Cerca et al., 2007).

Pre-Cretaceous Continental Metamorphic Block

The continental metamorphic core of southern Mexico was assembled during the Paleozoic, when the Mixteco and Oaxaca terranes were tectonically juxtaposed (Ortega-Gutiérrez, 1981; Ortega-Gutiérrez et al., 1999; Elías-Herrera and Ortega-Gutiérrez, 2002; Talavera-Mendoza et al. 2005), and they have shared a similar history of sedimentation and deformation since the Cretaceous (Nieto-Samaniego et al., 2006). The geometry of the shortening structures affecting the Mesozoic cover defines a broad, northward-convex arc showing an outward vergence away from the metamorphic core (Ferrari et al., 1998; Cerca et al., 2004; Nieto-Samaniego et al., 2006), suggesting rotation and/or vertical movements of the metamorphic block locally followed by gravitational detachment of the cover.


Model Setup

Modeling was performed at the Tectonic Modeling Laboratory of the Institute of Geosciences and Earth Resources (National Research Council of Italy) at the Earth-Sciences Department of the University of Florence. The simplified model was constructed considering an idealized initial strength profile for each region (e.g., Corti et al., 2003). We performed a series of four experiments in which two preexisting built-in lithosphere-scale weak zones (Arcelia–Palmar Chico and Guerrero Morelos Platform basins) were subject to successive orthogonal extension and oblique shortening (Fig. 4). The basins were simulated by stripes with a weaker strength profile (three-layer) welded within the normal strength profile (four-layer; Fig. 4). The mechanically stratified lithosphere was constructed with alternating layers of quartz sand and silicone + sand mixtures with properly scaled density (Table 1). Isostatic equilibration during deformation was achieved through a dense glycerol-gypsum mixture, which simulates the asthenosphere upon which the model was floating.

Models were deformed in one phase of orthogonal extension (velocity of 7.5 mm h−1) followed by oblique shortening at an angle of 15° (velocity of 9 mm h−1), representing the Early Cretaceous extension followed by the Late Cretaceous shortening.

In both phases, the wall was displaced by 30 mm (around 18% of extension and shortening, respectively). During the extensional phase, two metallic perpendicular plates provoked a velocity discontinuity (VD) in the middle part of the Arcelia–Palmar Chico basin (Fig. 4). One model (Laramide 01) consisted of a single phase of extension; the other models (Laramide 02–04) were subsequently shortened. In model Laramide 02, extension and compression were orthogonal, whereas the last two models were deformed obliquely at an angle of 15°. The angle of convergence between the Pacific and North American plates was calculated by Schaaf et al. (1995) for the Laramide time. Before starting the shortening phase, the mobile wall was rotated counterclockwise, and the gap was filled with material similar in strength to the normal four-layer lithosphere. In two models (Laramide 02 and 03), black sand was added to fill the basins formed during the extensional phase to simulate syntectonic sedimentation.

Top-view pictures of the model were obtained at regular time intervals. At the end of the experiment, models were soaked in water, frozen, and sectioned in longitudinal stripes to document and photograph cross sections.

Model Simplifications, Selection of Analogue Materials, and Properties

Analogue models allow us to integrate conceptual models of geological evolution in concrete terms but necessarily simplify rheological and/or geometrical parameters that exert a potential control on resulting deformation. In the first place, our models are not a realistic representation of the lithosphere, but an idealization inferred from geological information. The use of homogeneous materials with depth-invariant properties for representing the strongly temperature-dependent rheologies of the lower crust and ductile mantle is a major simplification employed commonly in similar experiments (Davy and Cobbold, 1991; Burg et al., 2002; Sokoutis et al., 2000). The sharp lateral contrast of strength along a vertical contact is also a simplification in our models. Lateral contrast in rheology is likely to produce localized instabilities, secondary flow in viscous layers (e.g., Montési and Zuber, 2003), and disturbing buckling wavelength and amplitude.

In our setup design, the moving wall transmitted lateral tectonic driving forces. In the case of extension, the experimental design is intended to simulate the hypothetical conditions before shortening and does not consider complexities derived from nonorthogonal extension. Several other factors having a great influence in the result of deformation were not considered, such as erosion, the effect of pore pressure in the generation and growth of structures, thermal evolution, and addition of new material in the crust by magmatism. Sedimentation was considered in three of the experiments in which the extensional basins were filled before the shortening phase.

Our model aims to simulate the structural style due to mechanical heterogeneity of the lithosphere and does not consider the thermal evolution. Nevertheless, a warm geothermal gradient below the basins can be the cause of the weak zones in the lithosphere (e.g., Boutelier et al., 2003), which allow and localize the deformation of an orogen during shortening (e.g., Chapman and Furlong, 1992; Pollack et al., 1993; Hyndman et al., 2005; Hartz and Podlachikov, 2008).

Model Scaling Analysis

The model generically explores the influence of lateral mechanical variations during successive phases of extension and compression, and it focuses on the structural style that resulted in the upper crust for the comparison with the natural prototype. The model is a scale representation of the prototype under study following within the possible principles of geometric, dynamic, and rheological scaling (Hubert, 1937; Ramberg, 1981; Weijermars and Schmeling, 1986). In this case, we used a geometrical scale such that 1 cm in the model represents 40 km in nature, l* = lmodel/lnature = 2.5 × 10−7. Stress scaling is obtained by applying the general equation for reducing gravitational stress (σ* = ρ*g*l*; where * denotes the ratio model/nature of each parameter, g* = 1 in the case of normal condition of gravity, and l* is the geometrical scale). The ratio of gravitational to differential stress in the viscous layers (Ramberg number; Weijermars and Schmeling, 1986) provides a measure for testing dynamic scaling, and it is given by
where ρ and hd are the density and thickness of the ductile layer, respectively, g is the acceleration due to gravity, and (σ1 – σ3)viscous is the differential stress that can be expressed also as (σ1 – σ3)viscous = ηė, where η is the viscosity of the viscous layer and ė is the strain rate given in the case of shortening by the ratio of the mean velocity of convergence, V, and the thickness of the ductile layer hd. Table 2 documents the scaling procedure for upper brittle and ductile layers considering the initial strength profile at the onset of the second phase of shortening.


In this section, the proximal side to the moving wall is referred to as hinterland, whereas the distal side corresponds to the foreland. The cratonic area is located to the east in the distal position, and the north of the models is located at the upper part of photographs. The model blocks were given the names of the prototype blocks or basins: Guerrero Morelos Platform (GMP or basin 1, B1) and Arcelia–Palmar Chico (APC or basin 2, B2) for the basins and Tejupilco anticlinorium (TA) for the area between the basins. The results are presented in two parts that relate to two main issues addressed in this work: (1) the surface evolution of deformational patterns during oblique shortening, and (2) the final result in vertical section of the models.

Deformation of the Two Phases in Top View

Extension Phase

The models share a similar top-view evolution exemplified by pictures taken during the Laramide 04 experiment shown in Figure 5A. The evolution of structures indicates a general westward migration of extension. The first major normal fault follows the geometry of the cratonic area, with a general N-S orientation that turns toward the NE in its northern part, and that formed in the models at the western margin of the cratonic area (at ∼5% of extension, not shown in the figure). Farther north in the weak zone, the fault acquires again a N-S orientation. Subsidence activated in the middle part of the Guerrero Morelos Platform (basin 1), and normal faults began to propagate from the lateral fixed walls toward the center of the model. By 6% of extension, a normal fault formed in the western margin of the Arcelia–Palmar Chico (basin 2). Concentration of extension in the western margin of the cratonic area implies that mechanical contrast with the weak zone amplifies deformation (e.g., Bonini et al., 2007). The Tejupilco anticlinorium is an uplifted area between both basins. By 8% of extension, normal faults with arcuate surface traces formed west of the cratonic area in the northern part of the model, and a second strip of extensional faults localized above the eastern side of the Tejupilco anticlinorium. After 11% of extension, deformation localized in the middle of basin 2, rapidly developing a set of rift-oblique faults between the two normal faults delimiting the rift center. Figure 5B shows a summary of the evolution of extension in the basins for two experiments (Laramide 02 and 04). In this figure, the width of each basin (Wbasin) along a line located at y ∼9 cm from the south for models Laramide 02 and 04 was divided by the total initial width of the model (Wi = 170 mm) and plotted versus the amount of deformation for consecutive increments of extension. Although the velocity discontinuity is located in basin 2, normal faults are observed at the surface first in the eastern margin of basin 1; after only 5% of extension, faulting propagates into basin 2. After 6% of extension in both experiments, the width of basin 1 becomes relatively stable with continuous subsidence. On the other hand, basin 2 widens linearly up to values of Wbasin/Wi of 0.35 at 18% of extension.

Oblique Shortening Phase

The evolution of oblique shortening is exemplified in Figure 6A. In these experiments, bulk shortening refers to the advance of the moving wall divided by the initial length of the model. In this case, shortening deformation reactivates all the previously formed structures as thrust faults. Thrusting in the eastern boundary of B2 and in both boundaries of B1 was observed after 7.5% of bulk shortening in the southern part of the model, and uplifting of both basins becomes evident. Reactivation of normal faulting is observed north of the cratonic block in B1. After 9% of bulk shortening, foreland advance of the reactivated normal faults occurs on the eastern margin of B2, and there is gentle uplift of the horst block (Tejupilco anticlinorium).

The relative width of the basins with respect to their original widths is plotted in Figure 6B for the Laramide 02 experiment, in which shortening was parallel and opposite to extension. Both basins show a similar decrease in width characterized by episodes of major decrease between 3.5% and 5% of bulk shortening, and after 9% bulk shortening, respectively. The first width decrease seems to be related to strength loss caused by folding of the ductile layers in the weak areas and early stages of reactivation and folding of the sand layer filling the basins, whereas the second episode is related to a major uplift and advance of reactivated thrusting. Comparison of final top-view photographs of the extension and shortening phases (Figs. 5 and 6) suggests that shortening deformation was accommodated preferentially by folding and reactivation of preexisting faults.

Model Deformation in Cross Section

The transverse sections of experiments provide a detailed view of the vertical accommodation of deformation in models (Fig. 7). During extension, deformation of ductile layers in the models is characterized by thinning in the weaker crust area. The lateral rheological boundaries with the normal and cratonic crust control the location and propagation of normal faults in the upper brittle crust (Fig. 7A).

After shortening, the final deformation pattern of the inverted extensional basins was controlled by the geometry of early normal faults, confirming previous observations in analogue brittle-ductile systems (Fig. 7B; e.g., Brun and Nalpas, 1996; Dubois et al., 2002; Del Ventisette et al., 2006). The mechanical heterogeneity of the blocks in the models and the presence of the preexisting normal faults result in shortening structures that do not follow a linear sequential and progressive development. For instance, a greater strength contrast of the weaker crust with the cratonic area resulted in an inverted basin with two opposite vergences. Figure 8 shows transverse sections of experiment Laramide 04 that cut at different positions from the lateral fixed walls. In section “a,” deformation is distributed along the entire section, but it can be observed that inversion of the brittle crust faults and relief uplift took place mainly in the weak zones. From top-view photographs, the easternmost fault is activated at early stages of shortening (∼7.5% of bulk shortening). The response of the lithospheric system is characterized by gentle anticline folding of the weak crust section. From section “a,” trough “c,” it can be seen that a large-scale anticline is formed by the uplifting of the contact between weak and normal crusts near the moving wall. The flank formed by the normal crust deforms mainly by folding, with different kinematics separated by the brittle mantle. Indentation of the stiffer lower crust between the upper and weak lower crust is favored by the relative strength contrast forming a syncline; in contrast, the lower ductile mantle is thrust by the weaker mantle, producing an anticline. During progressive shortening, the weak crust accommodates strain by folding, but mainly by thickening and extrusion (Fig. 8D).


Geological Constraints and Assumptions to Construct a Conceptual Model for the Inversion Tectonics in Southern Mexico

Several authors have suggested, explicitly or implicitly, the idea of inversion tectonics in southern Mexico for explaining local deformation features observed during Late Cretaceous shortening (Campa-Uranga and Ramírez, 1979; Cabral-Cano et al., 2000; Elías-Herrera et al., 2000; Elías-Herrera, 2004; Centeno-García et al., 2008; Cerca et al., 2007; Martini et al., 2009). In this work, we integrate the geological observations to produce a model of inversion tectonics to explain the structural styles observed in southern Mexico. The experimental results suggest that inversion tectonics are a plausible model for explaining the diversity of structures found in southern Mexico; however, testing of this model would require independent data, such as a more complete set of field evidence and geophysical information. In order to constrain the model, we based our setup on the only available information, which essentially stems from surface geology.

The initial setup is based on the additional following considerations:

1. Continental-scale convergence with eastward subduction of oceanic plates below the North American plate has been active at least since the Early Cretaceous (130 Ma), as suggested by continental-scale seismic tomography and numerical modeling (Grand et al., 1997; Lithgow-Bertelloni and Richards, 1998; Bunge and Grand, 2000), leaving no space for models postulating accretion of intraoceanic arcs. Our model considers an already assembled continental lithosphere at least since the Early Cretaceous.

2. The Early Cretaceous extensional phase geometry and kinematics are practically unknown and are inferred here from geological information. We propose that the onset of extension might have been marked by the deposition of the Zicapa conglomerate to the west of the continental margin in the Early Cretaceous.

3. Inversion resulted from the switch from extension to oblique shortening with main vergence toward the E-NE between the Turonian and the Paleocene. The inversion is localized and more evident in the Arcelia–Palmar Chico basin, in which volcanic rocks intercalated with pelagic limestone are thrust over the shallower Amatepec limestone (both limestone sequences have approximately the same age) and over the Tejupilco schist, and in the Guerrero Morelos Platform.

Comparison of Models with the Geological Data of Southern Mexico

A direct comparison between structures in the models and those of southern Mexico would overinterpret the available information (e.g., Burg et al., 2002); however, a discussion of the geological evidences and their comparison with the model results can help us to understand the mechanism of inversion tectonics. Particularly, the first-order characteristics of deformation, such as the propagation in space and time of deformation and the overall geometry of the surface structures, can be analyzed in relation to the crustal mechanical heterogeneity and the polyphase deformation that occurred during the tectonic evolution of the study area. In essence, modeling results show that basin inversion is a plausible mechanism to explain the diversity of structures observed in southern Mexico (Fig. 9).

Some mechanical results of modeling that can be applied in the comparison with the natural case include:

The inversion tectonics in models are largely influenced by the presence of a mechanical contrast in the lithosphere that can be produced by either long-lived preexisting discontinuities or faults produced in a previous phase of strain. The extensional phases produce a tectonic fabric in the crust that controls the geometry of shortening structures. With the low angle (15°) of oblique convergence, shortening in the brittle crust is accommodated mainly by preexisting extensional structures, and no new structures parallel to the mobile wall were formed (consistent with the modeling results of Brun and Nalpas, 1996; Dubois et al., 2002; Del Ventisette et al., 2006).

The inversion of basins is observed to occur as a response not only to reactivation of the preexisting normal faults, but also to thickening and continuous deformation of the ductile layers. The development of preferential vergence is toward the east in the model, but some sections showed an opposite vergence (Figs. 8C) that developed in a late stage of deformation, as documented in southern Mexico in the Tzitzio anticline (Martini et al., 2009). Eastward vergence was favored by the high strength contrast between the cratonic crust and the weak crust. In the model shown in Figure 9, the opposite vergence developed in the absence of the relatively stronger cratonic crust and with an initially horizontal contact between upper and lower crust. In the close-up areas of Figure 9, shortening in the upper crust can be observed to be accommodated by reactivation of preexisting normal faults and transport of the extended zones above a detachment surface. Major reverse faults (black thick solid lines in Fig. 9G) show normally a shallower dip in their upper part, probably formed by fault branches following a lower angle trajectory to the surface (McClay, 1989; Coward, 1994; Bump, 2003). The overall geometry of the east-vergent model orogen displays a similar geometry, since most of the abandoned faults retain their original normal displacements, whereas the active faults have reverse displacements. In the case of the Laramide 02 experiment, contraction of the lithosphere system resulted in thickening of the weaker zones and the formation of a large anticline corresponding to the Tejupilco anticlinorium.

The differential uplift of the contact between normal and weak crust in models (Fig. 8) suggests a lateral flow of asthenosphere and ductile crust toward the NE in order to accommodate deformation. Lateral variations in uplifting of the lithosphere were not observed in the model Laramide 02, in which extension and shortening were parallel, confirming that lateral flow of mass can contribute significantly to the final result of deformation: in this case, a continental root developed in the area with major relief. The results of modeling suggest that during oblique convergence, lateral mass transfer is important to maintain the equilibrium of the lithosphere, supporting the idea that some features can be explained by orogenic floating of the upper brittle crust above a ductile lithosphere trying to reach mass balance (Oldow et al., 1990) (Fig. 8C).

In summary, the roles of heterogeneity and deformation style of the lithospheric mantle in continental tectonics play a fundamental role in the accommodation of deformation and in reaching isostatic equilibrium in the models. The resulting orogen is far from being sequential and foreland progressive, and thus our models are capable of explaining the diversity of structural styles during a prolonged shortening phase. Our experimental setup implies that the mechanical heterogeneity was already established in the continental lithosphere in the Early Cretaceous, and the results suggest that this heterogeneity may have played a fundamental role in controlling the differences in geological history.


Simplified analogue models provide a first-order approximation for simulating intraplate deformation of southern Mexico lithosphere. The experiments allowed us to develop a parsimonious model explaining some aspects of Late Cretaceous shortening deformation, which are controlled by the previous extension and geological history. Based on the model results, we propose that the differences in geological history are controlled by the extension and shortening phases that affected the fragmented lithosphere of southern Mexico.

The features observed in the nature and explained by our model include:

(1) In models, as in nature, the shortening deformation front migrates eastward.

(2) Concentration of deformation at basin boundaries and the presence of opposite vergences are explained by the mechanical contrast between weaker and stronger crustal zones.

(3) Out of sequence structures and timing of shortening occur in different zones of the hinterland in the model and in nature. Different degrees of basement involvement (thick-skinned tectonics) and décollement (thin-skinned tectonics) were observed during the shortening phase, especially in its late stage, when folding of the lithosphere seems to be an important process controlling deformation, particularly for the presence of the uplifted Tejupilco block.

The models’ results confirm that previous extensional features and, more generally, lithospheric-scale discontinuities or strength contrasts are effective in transferring deformation over long distances. In the studied area, strain was resolved by reactivation of preexisting extensional faults. Predictions of the model may be tested and corroborated when more information about the structure of the southern Mexico lithosphere becomes available from geophysical studies.

The project was funded by grants CONACYT 42642 (to Ferrari) and 46235 (to Cerca) and PAPIIT (Universidad Nacional Autónoma de México) IN120305. Cerca acknowledges support of a bilateral project SRE (Mexico)–MAE (Italy) for a short stay in the modeling laboratory of Florence.