297699780-Oa

Research Paper

GEOSPHERE Interactions between propagating rifts and linear weaknesses in the lower crust

GEOSPHERE, v. 15, no. 5 Nicolas E. Molnar, Alexander R. Cruden, and Peter G. Betts School of Earth, Atmosphere and Environment, Monash University, Clayton, VIC 3800, Australia

https://doi.org/10.1130/GES02119.1

14 figures; 2 tables; 1 set of supplemental files ■■ ABSTRACT 1987; Smith and Mosley, 1993; Vauchez et al., 1998; Will and Frimmel, 2017). The nature of the rheological heterogeneities and/or mechanical anisotropies, CORRESPONDENCE: [email protected] Pre-existing structures in the crust such as shear zones, faults, and mobile whether discrete or pervasive, has a first-order influence on the distribution of belts are known to exert a significant control on the structural evolution of deformation (Morley, 1999). Many of these pre-existing features lead to linear CITATION: Molnar, N.E., Cruden, A.R., and Betts, P.G., 2019, Interactions between propagating rifts and lin‑ continental rifts. However, the influence of such features when the extension zones of anomalously weak material within the lithosphere. Heterogeneities ear weaknesses in the lower crust: Geosphere, v. 15, direction progressively changes over time remains uncertain. Here we present and anisotropies can therefore be treated as equivalents and are referred to no. 5, p. 1617–1640, https://doi.org/10.1130/GES02119.1. new results from three-dimensional lithospheric-scale laboratory experiments as weak seeds that can be incorporated in analogue models (e.g., Zwaan and of rotational extension that provide key insights into the temporal evolution Schreurs, 2017; Molnar et al., 2017, 2018). Factors that play a major role on rift Science Editor: Andrea Hampel Associate Editor: Valerio Acocella of propagating rifts. We specifically test and characterize how rifts propagate evolution include the location and obliquity of pre-existing structures with and interact with linear crustal rheological heterogeneities oriented at variable respect to the direction of extension (e.g., Acocella et al., 1999; Ebinger et al., Received 4 February 2019 angles with respect to the extension direction. Results show that approxi- 2000; Agostini et al., 2009) and temporal variations of the regional stress field Revision received 15 May 2019 mately rift-parallel pre-existing heterogeneities favor the formation of long, (e.g., Ebinger et al., 2000; Morley, 2010). Observations from many continental Accepted 18 July 2019 linear faults that reach near-final lengths at early stages. Low angles between rifts have greatly improved our understanding of rift evolution as a function of the heterogeneities and the propagating rift axis may result in strong strike- the obliquity of pre-existing weaknesses (e.g., Ring, 1994; Morley, 1999; Corti Published online 29 August 2019 slip reactivation of the pre-existing structures if they are suitably oriented et al., 2003; Morley et al., 2004; Autin et al., 2010). Natural examples show a with respect to the stretching direction. When the linear heterogeneities are wide variety of geometric relationships, covering the entire range between oriented at intermediate to high angles rift branches become laterally offset rift-subparallel weaknesses (e.g., Rukwa Rift, East Africa; Morley, 2010), to as they propagate, resulting in complex rhombic fault patterns. Rift-perpendic- approximately rift-perpendicular weaknesses (e.g., Main Ethiopian Rift, East ular crustal heterogeneities do not affect fault trends during rift propagation, Africa; Bonini et al., 2005, Corti et al., 2018a, 2018b) (Fig. 1). Understanding but cause stalling and deepening of laterally growing rift basins. Similarities the temporal evolution of rifts relies on the preservation of timing and/or kine- between the analogue experimental results and selected natural examples matic constraints within the geological record. When such evidence is lacking provide insights on how nature finds the preferential pathway to breakup in or when it is obscured by the superimposition of tectonic events, deciphering heterogeneous continental lithosphere. the temporal evolution of rifts can become a challenge. Complementary analogue (e.g., McClay and White, 1995; Acocella et al., 1999; Corti, 2008; Agostini et al., 2009; Erbello et al., 2016) and numerical exper- ■■ INTRODUCTION iments (e.g., Brune, 2014; Brune et al., 2017; Naliboff et al., 2017; Mondy et al., 2018) have provided important insights on the development and interaction The continental lithosphere is characterized by rheological heterogeneities of extensional faults, which help to constrain the overall temporal evolution and mechanical anisotropies inherited from previous tectonic events (e.g., of rifts. Temporal changes in the extension direction are a common feature of Audet and Bürgmann, 2011). Rheological heterogeneities are commonly asso- rifts in nature (e.g., Morley, 2010; Nemčok et al., 2016; Whittaker et al., 2016; ciated with past orogenic or intraplate processes (i.e., collisional belts, shear Pongwapee et al., 2018). Such changes have typically not been included in zones, rifts, thermal weakening) that result in areas of the lithosphere with geodynamic models of continental extension. Previous modeling studies that reduced strength (e.g., Vauchez et al., 1997; Corti, 2008; Gueydan et al., 2008). focused on the role of multi-phase rifts have successfully recreated different Mechanical anisotropies arise as a consequence of tectonic fabrics (i.e., foli- extension directions through time (e.g., Bonini et al., 1997; Keep and McClay, ations, lineations, crystallographic preferred orientations, reduced grain size) 1997; Henza et al., 2010; Erbello et al., 2016; Duffy et al., 2017), but achieved that define a mechanically weaker region than the surrounding lithosphere. It this by imposing laterally homogeneous extensional boundary conditions This paper is published under the terms of the has long been recognized that they control the pattern, propagation and overall during each phase. Conversely, the inclusion of a rotational kinematic boundary CC‑BY-NC license. structural and sedimentological evolution of continental rifts (e.g., Rosendahl, condition in experiments (e.g., Tron and Brun, 1991; Benes and Scott, 1996;

GEOSPHERE | Volume 15 | Number 5 Molnar et al. | Rift propagation across crustal weaknesses Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/15/5/1617/4831062/1617.pdf 1617 by guest on 27 July 2020 Research Paper

Crustal heterogeneity obliquity ( α ) 0˚ 15˚ 30˚ 45˚ 60˚ 75˚ 90˚

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Rukwa Rift Kenya Rift Gulf of Suez Reconcavo Graben Rio Grande Rift EARS W-branch

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50 km 50 km 20 km 50 km 70 km N Lake Tanganyika C Lake Tanganyika Gofa Province Upper Rhine Graben Malawi Rift

General extension Pre-existent weakness Major faults direction

Figure 1. Natural examples of continental rifts classified by the approximate angle of obliquity between the direction of extension and the controlling pre-existing crustal heterogeneities. Approxi- mate directions of heterogeneities from Morley et al., 1992 (Rukwa rift, East Africa), Versfelt and Rosendahl, 1989 (Lake Tanganyika, East Africa), Vétel et al., 2005 (Kenya Rift, East Africa), McClay and Khalil, 1998 (Gulf of Suez, northern Red Sea), Philippon et al., 2014 (Gofa Province, Ethiopia), Milani and Davison, 1998 (Recôncavo Graben, Brazil), Misra and Mukherjee, 2015 (Upper Rhine Graben, Germany/France), Zwaan and Schreurs, 2017 (Rio Grande Rift, southwestern USA/northern Mexico), Dawson et al., 2018 (Malawi Rift, East Africa), Corti et al., 2007 (East African Rift System [EARS] W-branch). Digital elevation maps (GeoMapApp [http://www.geomapapp.org]; Ryan et al., 2009) and simplified line drawings of major extensional structures. Note that red arrows indicate the extension direction at the moment of rift formation, which may differ from present-day kinematics. C—central; N—northern; W—western.

Molnar et al., 2017) has effectively simulated progressive changes in extension resolution, we characterize the detailed structural evolution of continental direction over time. Such changes may induce a laterally heterogeneous strain rifts as a function of the obliquity of crustal heterogeneities. We compare our field that results in a gradient in the amount of extension along the rift axis results to selected natural examples in order to provide new insights on rift (e.g., Gofa Province of the Main Ethiopian Rift; Philippon et al., 2014). When migration and propagation across discrete pre-existing crustal heterogeneities. these conditions are met, rift branches usually propagate laterally toward a pole of rotation. How they propagate across transversal pre-existing crustal heterogeneities remains poorly understood. ■■ EXPERIMENTAL PROCEDURE Although both pre-existing heterogeneities and rotational extension are often identified as important factors in ancient and modern rifts, their combined Model Setup influence has rarely been tested by means of analogue or numerical modeling (e.g., Benes and Scott, 1996; Molnar et al., 2017). In this study, we present a All laboratory experiments presented here are composed of a three-layer, series of three-dimensional laboratory experiments addressing this topic. By brittle-ductile model lithosphere that floats isostatically on a low-viscosity measuring surface deformation in the experiments at high spatio-temporal model asthenosphere, contained within an acrylic tank (Fig. 2). All models

Weak Normal 0 0 Upper crust Brittle 20 8 Upper crust Lower crust Ductile Qz sand + c-sph 35 14 Lithospheric mantle Ductile 80 30 Lower crust Asthenospheric mantle PDMS [Km] [mm]

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LA E A NW B 0˚ C 30˚ Free side

Asthenospheric mantle D 45˚ E 60˚ F 90˚ NS + NaCl

Figure 2. Experimental setup and boundary conditions. (A) Top: Three-dimensional sketch illustrating how the brittle-ductile multilayer lithosphere is constructed and positioned within two U-shaped containing walls. A side-cut of the model upper crust drawing shows how the weak seeds are incorporated into the model lithosphere. Bottom: cartoon drawing of the rotational extensional apparatus. Blue arrow indicates the direction in which the linear actuator (LA) pulls the moving wall at a constant velocity. A hinge between the U-shaped walls, on the opposite side of the linear actuator rod, acts as a fixed vertical pivot, creating a horizontal gradient in the imposed strain field (red arrows). Fixed pivot is designated as the north for reference. Qz—quartz; c-sph—ceramic spheres; PDMS—polydimethylsiloxane; BPL—black plasticine; GB—glass bubbles; NS—Natrosol; NaCl—sodium chloride. (B) Cross section showing the rheological layering of the model lithosphere and representative strength profiles corresponding to the normal and weak sections of the crust (see text and Table 1 for details). (C) Schematic top view drawing showing the motion described by the opening plate (in gray) and the subsequent gradient in extension produced along the rift axis (red arrows). (D) Schematic top view drawing indicating the initial location of the heterogeneities and the definition of the angle of obliquity α( ). Model D (α = 45°) boundary conditions are used as an example. ED—extension direction; RP—general rift propagation direction. (E) Graphical summary of the boundary conditions for all models discussed in this study. Colored rectangles show the area analyzed and presented in Figures 3–8. Each color indicates the location of the proximal (red), central (green), and distal (blue) domains for each model, delimited by the distance to the pole of rotation (models A and B) or by the position of the linear heterogeneities (models C to F). NW—no weakness.

have initial dimensions of 44 × 44 × 3 cm and are constructed to sit within two corresponds to ~16 mm/yr in nature. Scaling factors and experimental param- U-shaped confining walls (Fig. 2A). The experimental setup was designed to eters are detailed in Table 1. simulate continental rifts that propagate and interact with variably oriented The model lithosphere is composed of a mixture of quartz sand and hollow crustal anisotropies. We therefore mimic propagating rifts by imposing a rota- ceramic microspheres to replicate the brittle upper crust, polydimethylsiloxane tional kinematic boundary condition on the models following the approach (PDMS) to model the ductile lower crust, and a mixture of black Colorific used in Molnar et al. (2017). The model lithospheric plate is attached to a mov- Plasticine®, 3M® hollow glass microspheres and PDMS to replicate the litho- ing U-shaped wall and pulled by a linear actuator at a constant divergence rate spheric mantle (Figs. 2A and 2B). The upper crustal granular material mixture (Fig. 2A). A hinge between the two confining walls acts as a fixed vertical pivot, obeys the Mohr-Coulomb failure criterion, with an angle of internal friction creating a horizontal gradient in the imposed extensional strain field (Fig. 2A). ϕ< 38° and cohesion ~9 Pa (Molnar et al., 2017). PDMS and the PDMS-based The amount of extension in the models increases away from the pivot (Fig. 2C), mixture used for the ductile layers are quasi-Newtonian fluids frequently used which is analogous to creating an along-rift axis gradient in the amount of in analogue modeling studies (e.g., Weijermars, 1986; Cruden et al., 2006; extension, as inferred from modern examples (e.g., Gofa Province; Philippon Riller et al., 2012; Molnar et al., 2017). The model lithosphere overlies a model et al., 2014; Woodlark Basin, southwestern Pacific Ocean; Mondy et al., 2018). asthenosphere that contains a solution of Natrosol® 250 HH plus sodium chloride in deionized water (Davaille, 1999; Boutelier et al., 2016). A complete description of the experimental setup, materials, scaling, and model construc- Materials, Scaling, and Rheological Layering tion can be found in Molnar et al. (2017). We introduced discrete pre-existing crustal anisotropies (Morley, 1999) into We scale down forces, length, and time using dimensional scaling theory the experiments following previous analogue modeling approaches (e.g., Corti, (Hubbert, 1937; Ramberg, 1967; Davy and Cobbold, 1991) in order to achieve 2008; Le Calvez and Vendeville, 2002; Zwaan et al., 2016). Two ~3-mm-wide an appropriately sized model that is kinematically and dynamically similar and ~1-mm-high linear semi-cylindrical seeds of PDMS were positioned on top to the natural system. The scaling ratios (Table 1) for the experiments were of the homogenous lower crust prior to sifting the brittle upper crustal gran- set so that 1 cm in the models corresponds to 25 km in nature and 1 h in the ular mixture. Since the brittle upper crust determines the maximum strength experiment represents ~0.8 m.y. in nature. Since all experiments were carried of the model lithosphere (Fig. 2B), the inclusion of these linear seeds locally out under the normal field of gravity (1 g), the gravitational scaling ratio is 1:1. reduces its thickness and hence the integrated resistance of the lithosphere to The velocity scaling ratio is derived from the time, length, gravity, and viscosity deformation. This technique may be used to simulate planar zones of weak- scales, such that the divergence rate of ~5 mm/hr used in the experiments ness such as fault planes, shear zones or areas of previously thinned crust

TABE 1. SCAING PARAMETERS FOR THE ANAOGE MODES Thickness Density Viscosity Material Model Nature Model Nature Model Nature mm km kg/m3 kg/m3 Pa s Pa s pper crust Brittle 8 20 928 2600 ‑ ‑ S ESPH ower crust Ductile 6 15 982 2760 4 104 2 1021 PDMS

inear weakness pper crust Brittle 7 17.5 928 2600 ‑ ‑ S ESPH ower crust Ductile 7 17.5 982 2760 4 104 2 1021 PDMS

ithospheric mantle Ductile 16 45 1090 3050 2 105 1 1022 PDMS BP 1 Asthenosphere ‑ ‑ 1100 3100 380 1.9 1019 NaCl‑NS Scaling factors: Model/prototype 4 10–7 ρ 0.355 –17 Time scaling factor η ρ. g. .t t 1.41 1010 1 hr in model 0.8 Ma in nature Velocity scaling factor v l/ t v 2.84 103 6 mm/hr in model 20 mm/yr in nature Gravity scaling factor g gm/gp 1 Notes: All layers have dimensions of 440 440 mm. Model asthenosphere area is 626 626 mm. Ssand; ESPHhollow ceramic spheres; PDMSpolydimethylsiloane; BPblack plasticine; 1hollow glass microspheres; NaCl‑NSsodium chlorideNatrosol solution.

(Morley, 1999; Zwaan et al., 2016), although the thickness of the weak seeds TABE 2. IST OF THE PERFORMED ANAOGE MODES (i.e., <10 km) would be more appropriate to compare with strike-slip fault zones Model Obliuity angle α eakness orientation or major ductile shear zones in crystalline basement that have been exhumed A No weakness to upper crustal levels. B 0 N/A Rather than trying to reproduce a specific natural scenario, our aim is to C 30 Clockwise characterize how rifts propagate across crustal weaknesses as a function of C2 30 Anticlockwise their obliquity (α) with respect to the initial extension-normal direction (Fig. 2D). D 45 Clockwise We set a distance of 5 cm between the two discrete linear anisotropies in order D2 45 Anticlockwise to allow enough space for rift branches to develop freely between them and E 60 Clockwise E2 60 Anticlockwise to study their propagation and temporal evolution. For ease of analysis and F 90 N/A discussion, we subdivide the resulting fault pattern into proximal, central, G 15 Clockwise and distal domains, which are delimited by the position of the linear aniso G2 15 Anticlockwise tropies (Fig. 2E). H 75 Clockwise H2 75 Anticlockwise Eperiments described in this work. N/Anot applicable Deformation Monitoring when the weakness is either 0 or 90, there is no clockwise or anticlockwise orientation of the weaknesses.

We employed a particle image velocimetry (PIV) system to monitor deformation and surface topography during the experiments. The PIV system consists of two stereoscopic cameras and one top camera that are set to capture digital the experimental time (X axis) during which each structure had been accom- images of the experiments at 2 min intervals, which provides quasi-continuous modating extension, arranged by the distance to the pole of rotation (Y axis). monitoring of deformation throughout the entire 36 h duration of the experi- In this way, it is possible to visualize the timing of activity/inactivity for the ments. The digital images are processed using a cross-correlation algorithm main extensional faults and depict the temporal evolution of the main rift (Adam et al., 2005), which allows the extraction of high-resolution surface faults for each of the previously defined proximal, central, and distal domains. topography, differential and cumulative strain, and velocity data from the models. See Supplemental Data1 for details on the methods used for PIV data calculations. Model A (Homogeneous Reference Lithosphere)

Strain initially localized in the distal domain as extension was accommo- ■■ RESULTS dated by linear, normal faults that formed perpendicular to the extension direction (Fig. 3A). Increasing stretching resulted in the propagation of these A total of 12 experiments were carried out. Each experiment had identical major boundary faults toward the pole of rotation, forming northward pointing rheological layering and kinematic boundary conditions. The angle of obliq- V-shaped basins (Figs. 3A–3B). Differential strain measurements indicate the uity of the crustal anisotropies was increased by 15° steps between each incipient formation of new rift graben in the central domain, bounded by linear experiment. We present and describe the detailed results of six experiments normal faults that also strike approximately perpendicular to the stretching in which the linear anisotropies were progressively rotated clockwise between direction (Fig. 3B). After 18–19 h (14.6–15.4 m.y.), activity on the distal domain experiments. Additional experiments with anticlockwise-rotated weak seeds progressively decreased, rift tip propagation slowed down and the central were carried out for comparison and showed similar patterns and timing domain rift graben started to accommodate most of the extension (Fig. 3D). of deformation. Figure 2E and Table 2 summarize the initial conditions of The bounding normal faults in the central domain formed V-shaped graben 1 Supplemental Materials. Detailed explanation of the the experiments. Results are divided into three stages based on experiment that propagated both north and south, with the eastern graben developing deformation and strain monitoring provided in the running time and the corresponding amount of displacement imposed at the more rapidly and forming deeper basins (Figs. 3B–3C). supplemental materials, including workflow and vi- free side of the model. The pole of rotation is designated as the north side The overall deformation was gradually transferred toward the east and to sual representation of differential displacement field maps calculations. Figures of analogue models not for reference; hence, the amount of extension increases from north to south the proximal domain, with faults that grew toward the pole of rotation due to discussed in detail in the manuscript are additional- (Fig. 2C). Surface topography and differential strain for each time-step are the anticlockwise opening of the moving plate (Figs. 3A–3C). The main bound- ly presented in these materials. Please visit https:// used to display the general evolution of propagating rifts in each experiment. ary faults in the distal domain were active for 26 h (i.e., 21.1 m.y.) (Fig. 3D), after doi.org/10.1130/GES02119.S1 or access the full-text article on www.gsapubs.org to view the Supplemen- Rose diagrams of fault distributions and fault activity graphs are presented which strain was transferred to the rift depressions (Fig. 3C). Normal faults tal Materials. for selected experiments. Fault activity graphs (panel D in Figs. 3–8) illustrate in the central domain accommodated most extension between 18 and 27 h

Time = 6.5 h Time = 16.5 h Time = 26.5 h 5.4 m.y. 13.5 m.y. 21.6 m.y. Extension = 8% Extension = 20% Extension = 32% 350NW 350 350 N N N

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Figure 3. Evolution of deformation for Model A with homogeneous lithosphere (i.e., no crustal anisotropy). Panels A, B, and C display top view model deformation at early, intermediate, and late stages, respectively. Top-view differential strain on surface and structural interpretation in white lines are shown in the left part of each panel. Top-view digital elevation models and structural interpretation in black lines are shown in the right part of each panel. Blue arrow indicates the general extension direction at each stage. White dashed lines indicate approximate location of the crustal heterogeneities. Rose diagrams at the bottom of each panel illustrate normalized fault distribution weighted for the fault length for the whole model (orange) and subdivided for each domain (red, green, and blue). Blue line in the rose diagrams indicate the general extension direction at each stage and the double red lines show the orientation of the crustal heterogeneities when present. (D) Left: Top view line drawing of structures interpreted at late stages of the experiment, colored according to the domain in which they formed. Right: Fault activity graph illustrating the timing of activity/inactivity for the main extensional faults. DiffS—differential data of normal strain on surface.

(i.e., 14.6 – 21.9 m.y.) (Fig. 3D). Finally, extension was predominantly accommo- central domains (Fig. 5A). The only structural features that developed out- dated in the proximal domain after 27 h (i.e., 21.9 m.y.). Comparison between side the linear weak area were two short, N-S–oriented faults that delimited fault orientations over the different time steps indicate that the boundary a rift depression in the distal domain of the model (Fig. 5A). With increasing faults were oriented perpendicular to the extension direction during the entire deformation, this rift branch propagated northward and interacted with the experiment, and rift graben propagated both away from and toward the pole southernmost anisotropy. This interaction resulted in a locally heterogeneous of rotation without being laterally offset as they propagated (Figs. 3A–3C). strain field with relatively high differential strain in the south of the anisotropy, while extension continued to be accommodated by the early-formed N30°E oriented elongated normal faults toward the north of the anisotropy (Fig. 5A). Model B (α = 0°) When deformation propagated beyond the anisotropy and into the central domain, the new rift bounding normal faults developed at high angles with Deformation became highly localized at early stages, forming two elon- respect to the stretching direction (Figs. 5B–5C). gated rifts aligned parallel to the crustal anisotropies. The presence of a linear Differential strain measurements indicate a progressive north-eastward weak seed produced anomalously long and straight bounding normal faults propagation of strain along the southernmost anisotropy (Figs. 5A–5C). Defor- that reach their near-final lengths early in the experiments (Fig. 4A). Differ- mation northeast of this anisotropy created a fault system with a horsetail ential strain measurements showed a more evolved stage in the eastern rift geometry (Fig. 5B), in which extension was accommodated by three curved graben, associated with its proximity to the moving plate (Fig. 4A). An incipient splays that propagated toward the pole of rotation (Figs. 5B–5C). The west- N-S–oriented V-shaped central rift branch, located between the two main rift ernmost rift branches of the horsetail splay were bounded by normal faults compartments, was detected by differential strain measurements before devel- trending approximately parallel to the central domain rift branch (Fig. 5C). oping a visible feature on the surface (Fig. 4A). Rift branches that propagated Strain was mainly localized along the southernmost crustal anisotropy northward beyond the crustal anisotropies showed a more irregular pattern, throughout the entire experiment (cf. Figs. 5A–5C). Consequently, only minor but orientation analysis indicates that the fault trends remained perpendicular sinistral oblique-slip extensional kinematics developed along the northern to the extension direction (Fig. 4A). crustal anisotropy and faults were completely absent in the proximal domain Deformation evolution followed the behavior observed at early stages, (Fig. 5C). This indicates that the low obliquity angle of the crustal anisotropy with the widening of the two main rift branches accommodating most of the favored dislocation of the model lithosphere into kinematically independent extension (Fig. 4B). Compared to the early stages, the differential strain field blocks, resulting in low relative strain within the central and proximal domains map shows maximum strain values closer to the pole of rotation, indicating (Figs. 5A–5C). a progressive northward migration of deformation (cf. Figs. 4A and 4B). Rifts beyond the northern tips of the linear anisotropies continued their northward propagation as short arcuate segments but with an overall N-S trend. Model D (α = 45°) Fault development in the proximal domain was not controlled by pre-existing anisotropies, resulting in smaller, slightly laterally offset rift basins that sub- Early stages of deformation were characterized by partitioning of defor- sequently merged into a single graben (Figs. 4A–4C). Although deformation mation into two N-S–oriented rift graben in the distal domain and two other in the eastern main branch initially propagated farther north, higher values of V-shaped rift basins bounded by N45°E oriented linear normal faults localized differential strain were recorded in the western main branch at advanced stages along the crustal anisotropies (Fig. 6A). These narrow structures have sinistral (cf. Figs. 4A and 4C). Secondary rift branches with faults trending perpendicular oblique-slip components, indicating that the weaker crust favored dislocation to the extension direction developed on both sides of the linear anisotropy of the model lithosphere into blocks with different displacement rates as in (Fig. 4C). During the final stages, maximum values of strain were observed Model C (α = 30°) (Figs. 6A–6C). in the proximal domain, near the pole of rotation, while deformation was Boundary faults of the distal domain rift branches propagated northward, more distributed and limited to the basin floors in the distal domain (Fig. 4C). gradually changing their orientation as they propagated toward the southernmost linear anisotropy (Fig. 6A). A series of three new rift branches with bounding faults trending ~N30°W formed in the central domain after 12 h Model C (α = 30°) (~9.7 m.y.) (Figs. 6B and 6D). These basins formed at slightly obtuse angles with respect to the extension direction, and at ~80° with respect to the weak- Deformation became strongly localized early in the experiment due to the ness-related faults (Figs. 6B and 6E). The orientation of the central domain presence of the N30°E oriented crustal anisotropies, as observed in Model B graben may be associated with a local modification of the stress field caused (α = 0°) (Fig. 5A). Extension was accommodated almost entirely by elongated, by horizontal displacement along the two linear anisotropies. This produced N30°E oriented sinistral oblique normal faults that dislocated the distal and a zig-zag pattern and consequently resulted in a rhombic basin geometry

Time = 6.5 h Time = 16.5 h Time = 26.5 h 5.4 m.y. 13.5 m.y. 21.6 m.y. Extension = 8% Extension = 20% Extension = 32% 3500° 350 350 N N N

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Figure 4. Evolution of deformation for Model B with obliquity angle α = 0°, illustrated as in Figure 3. Location of the crustal heterogeneities (Heter.) is illustrated as a dashed white line in panels A, B, and C. DiffS—differential data of normal strain on surface.

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Figure 5. Evolution of deformation for Model C with obliquity angle α = 30°, illustrated as in Figures 3 and 4. DiffS—differential data of normal strain on surface; Heter.—heterogeneities.

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Figure 6. Evolution of deformation for Model D with obliquity angle α = 45°, illustrated as in Figures 3 and 4. DiffS—differential data of normal strain on surface; Heter.—heterogeneities.

(Figs. 6B–6C). Rift branches that developed in the central domain were later- orientation of the bounding normal faults was not affected by the crustal ally offset with respect to those formed previously in the early formed distal anisotropies, but a pulsed northward progression of strain was detected in domain, except for the westernmost graben (Fig. 6B). the computed differential strain field maps (Fig. 8A). A decrease in the amount Differential strain measurements indicate that strain along the oblique-slip of strain toward the north coincided with the location of the rift-perpendicular fault zones propagated in a northeast direction (Figs. 6A–6C). Figure 6B shows linear anisotropies (Fig. 8A). This suggests that deformation migration stalled that strain along the northernmost oblique-slip fault zone stopped propagating in the proximity of the anisotropies and rift branches continued to propagate after ~16.5 h (~13.4 m.y.), when strain was relayed into the N30°W trending only after additional strain was accumulated (Figs. 8A–8B). proximal domain rifts (Fig. 6B). Rift branches in the proximal domain devel- Early in the experiment the eastern rift branch displayed a more advanced oped with a right-lateral offset with respect to those in the central domain deformation, having propagated beyond the southernmost linear weak area (Figs. 6B–6C). Orientation analysis of the proximal domain normal faults show and subsequently slowing down in the proximity of the northern anisotropy a progressive rotation toward the north due to their proximity to the pole of (Fig. 8A). However, with ongoing deformation, the western branch propagated rotation (Figs. 6C and 6E). Distal and central domain structures kept accommo- northward faster, and beyond the eastern branch (Figs. 8A–8C), showing a dating deformation during the final stages, but to a lesser extent than those in similar progression to Model B (cf. Figs. 4 and 8). the proximal domain (Figs. 6C and 6E). The highest differential strain values Maximum computed strain values were located closer to the northern end during the late stages were detected at the intersection between the northward of the model, indicating a progressive migration of deformation toward the propagating central domain graben and the northernmost crustal anisotropy. pole of rotation (Fig. 8C). Differential strain measurements at the final stages These areas, in turn, coincide with greater fault throws and hence formed the do not exhibit compartmentalization of deformation, indicating that the prop- deepest basins recorded in the experiment (Fig. 6C). agation-inhibiting role of the crustal anisotropies was mostly active during the incipient rupture of the crust and not at later stages of rift deepening and propagation (cf. Figs. 8A–8C). Nevertheless, digital elevation maps display an Model E (α = 60°) along-axis variation in rift basin depth, with deeper sections located between the crustal anisotropies (Fig. 8C). Extension at early stages was predominantly accommodated by N-S– trending faults that delimit two main rift branches that propagated northward, toward the pole of rotation (Fig. 7A). Differential strain measurements show Role of the Position and Orientation of Heterogeneities minor localization along the crustal anisotropies, which results in short, N60°E oriented sinistral oblique normal faults (Fig. 7A). The northward propagation A complementary set of experiments was carried out with an anticlockwise of the main rift branches was laterally deflected eastward by the presence orientation of the crustal heterogeneities and the same obliquity angles used of the anisotropies, as evidenced by the sinusoidal pattern of the differential above (models C2, D2, E2, G2, H2; Table 2; Fig. 9). Similar first-order struc- strain field (Fig. 7A). tural patterns and temporal evolution of deformation were observed between Most extension during the experiment was accommodated by N-S rift experiments with the same angle of obliquity but opposite anisotropy orien- branches (Figs. 7A–7C). Maximum computed values of differential strain at tation, with minor differences in the location of extensional faults outside the early stages were detected in the western rift branch, however, the faster anisotropic linear zones (Fig. 9). While the position and relative orientation northward propagation of the eastern rift indicates a progressive migration of of the heterogeneities with respect to the stretching direction has a signifi- deformation toward the moving plate (cf. Figs. 7A–7C). Advanced stages show cant impact, fault orientations and the influence of the crustal heterogeneities distributed deformation in the distal domain, with low strains localized in the on northward propagating rifts were comparable between experiments with rift depression, indicative of activity on internal rift faults or hyperextension in clockwise and anticlockwise oriented anisotropies. the ductile layers of the lithosphere (e.g., Manatschal, 2004) (Fig. 7C). Oblique- Comparison between differential displacement field maps of selected slip extensional faults associated with the crustal anisotropies decreased their clockwise (Models C, D, and E) and anticlockwise (Models C2, D2, and E2) activity with ongoing deformation and became completely inactive after 23 experiments provides insights on how the anisotropy orientation affects the h (~18.6 m.y.) (Fig. 7C). sense of motion on oblique-slip fault zones. Sinistral oblique-slip motion was detected early in the experiments when the linear anisotropies were oriented clockwise with respect to the rift propagation direction. Conversely, anticlock- Model F (α = 90°) wise oriented linear anisotropies favored dextral oblique-slip reactivation. The influence of the anisotropy orientation on the oblique-slip component of Two extension-normal rift branches developed early in the experiment motion explains the differences observed in the lateral offset of propagating and propagated northward defining narrow V-shaped basins (Fig. 8A). The rift branches between displaced domains (Fig. 9).

Time = 6.5 h Time = 16.5 h Time = 26.5 h 5.4 m.y. 13.5 m.y. 21.6 m.y. Extension = 8% Extension = 20% Extension = 32% 35060° 350 350 N N N

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al Distal Centr Distal

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Figure 7. Evolution of deformation for Model E with obliquity angle α = 60°, illustrated as in Figures 3 and 4. DiffS—differential data of normal strain on surface; Heter.—heterogeneities.

Time = 6.5 h Time = 16.5 h Time = 26.5 h 5.4 m.y. 13.5 m.y. 21.6 m.y. Extension = 8% Extension = 20% Extension = 32% 35090° 350 350 N N N

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Figure 8. Evolution of deformation for Model F with obliquity angle α = 90°, illustrated as in Figures 3 and 4. DiffS—differential data of normal strain on surface; Heter.—heterogeneities.

Crustal heterogeneity obliquity (α) 0˚ 15˚ 30˚ 45˚ 60˚ 75˚ 90˚

Model G Model C Model D Model E Model H Clockwise

Model B Model F

Model G2 Model C2 Model D2 Model E2 Model H2

Anticlockwise

Figure 9. Graphical summary of the analogue models arranged by the angle of obliquity of the crustal heterogeneities with respect to the general stretching direction, α. Line drawings of structures presented as an overlay to top-view digital elevation models. Comparison between clockwise and anticlockwise setting of the heterogeneities show close similarities in terms of fault pattern and final rift architecture.

■■ DISCUSSION This morphology resembles the classic wide rifting mode of extension (e.g., Benes and Davy, 1996) but with less distributed deformation. As rifts propa- Summary of Results gated toward the northern end, the deformed area narrowed (up to 100–125 km) and resulted in a style of deformation similar to a narrow rifting mode. The modeling results presented show how the presence and orientation of Differences in the spatio-temporal evolution and rift architecture between crustal anisotropies affects the propagation and evolution of continental rifts. experiments with similar rheological and kinematic initial conditions were The overall deformation pattern is broadly characterized by extensional struc- solely dependent on the angle of obliquity of the crustal anisotropies, α (Fig. 9). tures that initiate at the southern end of the model and propagate northward Extensional structures in models with α ≤ 30° (Models B, G, G2, C, and C2) were toward the pole of rotation, in agreement with the general structural evolution strongly localized along the crustal anisotropies at early stages. This localiza- previously described for propagating rifts (Courtillot, 1980; Hey et al., 1980; tion results from the high angle between the anisotropies and the maximum Courtillot, 1982, Martin, 1984). In the homogeneous lithosphere experiment, instantaneous stretching direction and persisted throughout the experiment, subparallel structures with a horst-graben style of deformation were distributed although Model B (α = 0°) showed a wider distribution of deformation at over a 10–11-cm-wide (250–275-km-wide) area in the south end of the model. advanced stages (Figs. 9 and 10). The overall fault development was greatly

controlled by the linear weak sections; long, linear normal faults that aligned Experiments with a linear anisotropy oriented between 30° and 60° (Models with the anisotropies accommodated most of the extension, while only minor D, D2, E, and E2; Fig. 11) were characterized by complex interactions between structures, trending approximately perpendicular to the extension direction, the northward propagating normal faults and the crustal anisotropies. Local- developed outside or beyond the linear weak areas. Displacement field evo- ization along the anisotropy was strongest at early stages, when normal fault lutionary maps (Figs. 10 and 11) illustrate how the reactivation of the linear sets with sinistral oblique-slip components developed, resulting in kinematic anisotropies as normal faults with a left-lateral component was easier when dislocation of the model lithosphere into discrete domains (Fig. 11). In both the obliquity angle was between 15° and 45°. This resulted in differential rela- experiments, the oblique-slip activity gradually decreased with ongoing tive displacement between the distal and central domains of the experiments, stretching and deformation was progressively transferred to extension-normal which was maximum for Model C (α = 30°, clockwise) (Fig. 10). structures. This interaction resulted in rift branches that are laterally offset

Time Time 6 h 18 h 30 h 6 h 18 h 30 h

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Figure 10. Differential displacement for models A to F over time. Shown field maps correspond to the displacement in a chosen direction,φ , as indicated in the bottom left panel for each model, which were calculated after decomposing absolute displacement values (see Supplemental Material [text footnote 1]). Colors represent the distance that the model surface has been displaced since the previous timeframe computed with the particle image velocimetry, which was arbitrarily defined as 20 min (i.e., equivalent to ~0.3 m.y.) for noise reduction and optimal visualization (see Supplemental Material for details on deformation monitoring). Line drawings of interpreted structures are shown as semi-transparent black lines for reference.

C 30˚C2 D 45˚ D2 E 60˚E2

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Figure 11. Clockwise versus anticlockwise comparison of differential displacement field maps for Models with α = 30° (C-C2), 45° (D-D2), and 60° (E-E2), computed and illustrated as in Figure 10. Values correspond 18 h 18 h 18 h to the displacement parallel to the orientation of the heterogeneities (α) in order to highlight the sense of motion of the ex- Time Time tensional structures. Small insets showing vector arrows are included to exemplify the relative amount of displacement between the different lithospheric blocks along the fault plane, from which the oblique-slip sense of motion along the crustal heterogeneities is inferred. 30 h 30 h 30 h

0 1 0 0.5 0 1 0 1 0 2 Differential displacement [mm] 10 mm

as they propagated northward across the anisotropies, resulting in rhombic Natural Examples fault-bounded domains (Figs. 9–11). When α > 60° (Models H, H2, and F), the acute angle between the anisot- The experimental results can be compared to selected natural examples of ropies and the stretching direction inhibited strain localization and extension continental rifts that have been affected by heterogeneous extensional stress was almost entirely accommodated by N-S–trending normal faults. The fields and/or that experienced changes in extension direction over time. Due crustal anisotropies had minimal effects on fault trends during their north- to the intrinsic complexity of natural rift systems, we cannot replicate all of ward propagation and induced minor lateral offsets (Figs. 9 and 10). A pulsed the variables that control the evolution of continental rifts. In addition, the propagation of the extensional basins was observed only when the crustal dimension of the deformed regions in the selected examples is lower than that heterogeneities were oriented parallel to the initial extension direction (Model observed in the experiments. However, we argue that this does not change F, Fig. 8). Deformation analyses show that strain accumulated in steps as the extrapolation of the main findings of modeling to nature and that our it propagated along the main rift branches, and only penetrated through experiments provide first-order insights on the spatio-temporal evolution of the pre-existing heterogeneities after a threshold strain was reached. This propagating rifts. mechanism of propagation consequently led to deeper rifts in the vicinity of Here, we discuss comparisons between our experimental results and natu- the crustal anisotropies. ral examples as a function of the angle of obliquity of the crustal anisotropies

with respect to the extension-normal direction, α. For ease of visualization, Barth, 2010; Saria et al., 2014), but the timing of events that led to the formation some modeling results are mirrored and/or rotated in Figure 12 to compare of these NW-SE–oriented narrow basins remains controversial (Morley, 2010 with the natural prototypes. and references therein). It is widely accepted that a series of NW-SE rift basins of Karoo age (Permo–Triassic) (Morley et al., 1992; Morley, 1999; Delvaux, 2001) played an important role in controlling the Oligocene to present-day (Roberts Low Obliquity (α ≤ 30°) et al., 2012) rifting of the Tanganyika-Rukwa area. The general direction of extension affecting the EARS is a contentious issue (e.g., Ebinger, 1989; Ring Tanganyika-Rukwa Rift Segments et al., 1992; Morley, 1999; Chorowicz, 2005; Calais et al., 2006; Bellahsen et al., 2013) (Fig. 12A), but recent studies support a general E-W regional present day The Tanganyika and Rukwa rift basins are located in the central segment extension (Corti et al., 2007; Delvaux and Barth, 2010; Heidbach et al., 2010; of the Western Branch of the East African Rift System (EARS) (Fig. 12A). Pres- Morley, 2010). Additionally, two-dimensional numerical simulations indicate ent-day kinematics and recent fault activity is well constrained by geodetic, that larger rifts may capture rigid cratons, which can lead to local rotations geophysical, and focal mechanism data (e.g., Morley et al., 1992; Delvaux and (Koehn et al., 2008). These rotations accommodate the orthogonal motion of

N N P 6˚N Eastern 3.2 Nubian branch Plate 5˚S Victoria 0˚ Western Plate Arcuate branch 3.7

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300 km b LM UR Linear 3.8 9˚S 28˚E 34˚E 40˚E

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mm 100 mm Models C2 / G2 -4

Figure 12. Comparison with nature: Tanganyika-Rukwa rift segments (East Africa). (A) Cartoon illustration of the regional geological setting (top) and comparison with low obliquity models (α ≤ 30°) (bottom) Blue arrow indicates the sense of rotation of the moving plate (in gray). Red arrows illustrate the regional stretching direction. Model results are rotated 15° for comparison. (B) Structural map of the Tanganyika-Rukwa rifts (after Morley et al., 1992). Orange arrows illustrate the inferred local stretching direction as supported by focal mechanisms (Morley, 2010) and purple arrows indicate Victoria-Nubia plates relative motion in mm/yr (Stamps et al., 2008). Red dashed lines indicate the location of the inferred pre-existing crustal weaknesses, inherited from a Karroo-age extensional tectonic stage. LM—Lake Mweru; UR:—Usangu rift. (C) Top view digital elevation models of Model B and G2 showing the change in fault style from linear faults along the anisotropic area to short, arcuate faults beyond them. (D) Ex- ample of a propagating rift branch ignoring a pre-existing heterogeneity, as interpreted in Central Lake Tanganyika. (E) Subsidence/uplift maps of the timeframe 18–28 h for models G and G2. KA—Karoo Age rift basins; AS—Aswa shear zone.

larger plates and are consistent with the present-day anticlockwise rotation of two heterogeneity-related rift branches accommodated most of the extension the Victoria plate—containing the Tanzania craton—with respect to the Nubian early in the experiments, but soon after deformation was transferred to a single plate (Calais et al., 2006; Koehn et al., 2008; Stamps et al., 2008) (Fig. 12A). rift branch with larger fault throw and therefore more accommodation space for The Tanganyika and Rukwa rifts are good analogues for the experiments with sedimentation (Fig. 12C). The experimental evolution of deformation also indi- low obliquity due to the characteristics of the boundary conditions; the Karoo- cates strong uplift of the western rift shoulders, which compares to the higher age pre-existing discontinuities are oriented almost perpendicular to the local elevations in the western areas of Lake Tanganyika and Lake Rukwa (cf. Figs. 12B direction of extension and there is a comparable rotational component in the and 12E), which were probably aided by magma underplating (Morley et al., 1992). extensional system Fig. 12 and Supplemental Figs. S2–S4 [footnote 1]). A series of Pleistocene NE-SW–trending extensional structures (e.g., The Tanganyika-Rukwa rifts are narrow, deep basins bounded by a series Usangu rift and Lake Mweru) formed perpendicular to the main boundary faults of anomalously long, straight NW-SE–trending faults that follow pre-existing of the Tanganyika-Rukwa rifts (Ebinger, 1989; Morley et al., 1992) (Fig. 12B). heterogeneities (Morley et al., 1992; Ring, 1994; Corti et al., 2007) (Fig. 12B). Field observations also suggest a relative change in the extension direction that Pure dip-slip motion characterized these long, straight boundary faults at reactivated earlier dip-slip faults as strike- and oblique-slip faults (Ring, 1994; early stages (Foster and Jackson, 1998; Delvaux, 2001; Morley, 2010), indicat- Morley, 1999). The absence of normal faults with intermediate trends between ing a local stretching sub-perpendicular to the pre-existing heterogeneities the Tanganyika-Rukwa rifts and the Pleistocene NE-SW–oriented extensional (Fig. 12B). Since the regional extension direction was roughly E-W, this local structures indicate a rapid change in the extension direction. Although this feature could be explained by a deflection of stresses by inherently weak change is supported by field observations, it may be only a local feature basement fabrics (Morley, 2010). A similar behavior was observed in our low since some plate kinematic models (e.g., DeMets and Mekouriev, 2016) can- obliquity experiments, which developed long, linear, dip-slip normal faults not explain a variation in the regional extension direction. Models B, C, C2, (Fig. 9 and Supplemental Figs. S2-S4). The overall rift pattern in these models G, and G2 show a history of rift propagation that is partially analogous to the was reached at early stages and remained almost constant throughout the Oligocene–Pleistocene stages of the Tanganyika-Rukwa rifts. However, they do whole experiment (Figs. 4 and 5). This is also consistent with field observations not explain the perpendicular structures that developed later, which implies a in the Tunganyika-Rukwa rifts, which suggests that deformation pattern was significant change in the extension direction that did not occur in our exper- relatively static since its initial formation (Morley, 2002). iments. Such an abrupt change also explains dextral strike-slip reactivation The deflection of the Tanganyika rift from a NW-SE direction to a more N-S of the main NW-SE–trending faults (Ring, 1994). Since strain tends to keep orientation can be explained by the closer proximity of the northward prop- accumulating in the same locations (Morley, 2002), the pre-Pleistocene normal agating rift to the relative pole of rotation of the Victoria plate (Stamps et al., faults acted as pre-existing heterogeneities that facilitated their reactivation 2008) (Fig. 12A). All of our low obliquity experiments showed strong initial under the new stretching regime. At the same time, our modeling results localization along pre-existing heterogeneities, followed by later deflection of show that the left-lateral motion between domains was strongest when the propagating rifts toward the pole of rotation (Figs. 4, 5, and 9). Models B, C2, G, weaknesses where favorably oriented (i.e., 30°–45° clockwise, Fig. 11), and is and G2 have partial resemblance to this segment of the EAR. Only Lake Rukwa consistent with a ~60° to 90° change in the extension direction (Versfelt and and the southern section of Lake Tanganyika were strongly influenced by pre-ex- Rosendahl, 1989; Delvaux et al., 1992; Ring et al., 1992) (Fig. 12B). isting heterogeneities (Morley, 1999, 2010), which favored the formation of long linear faults near the anisotropy but led to more sigmoidal or arcuate normal faults in the central and northern Lake Tanganyika areas (Fig. 12B). Similar Intermediate Obliquity (30° < α ≤ 60°) observations were replicated in Models B and G2 (Fig. 12C). The development of these sigmoidal faults probably occurred synchronously with the northward Recôncavo-Tucano-Jatobá Basins propagation of the rift (Macgregor et al., 2017). Smaller scale discrete fabrics in the upper crust also contribute to this pattern, as shown by previous analogue The Recôncavo-Tucano-Jatobá rift basins of northeastern Brazil formed models of fabric reactivation (Corti et al., 2007). Additionally, an area in which during the Neocomian during the opening of the South Atlantic Ocean (Milani a propagating rift branch ignored a pre-existing heterogeneity was observed in and Davison, 1988; Heine et al., 2013). Previous studies (e.g., Milani and Davi- Model C2 (Fig. 12D), which could be analogous to the region west of southern son, 1988; Darros de Matos, 1992, 2000) have shown that complex structural Lake Tanganyika, in which no significant reactivation occurred. framework of these rift basins was strongly controlled by pre-existing crustal Sediment thicknesses in the Rukwa rift are less than in the more extension-or- heterogeneities associated with E-W to NE-SW–trending Pan-African shear thogonal central Tanganyika rift (Chorowicz, 2005; Corti et al., 2007; Misra and zones (Milani, 1987; Magnavita, 1992; Fig. 13A). Several interpretations have Mukherjee, 2015) (Fig. 12B). This is similar to the evolution of models G and proposed an anticlockwise rotation of the East Brazilian microplate (e.g., Milani G2, in which the extensional basins that formed perpendicular to the stretching and Davison, 1988; Szatmari and Milani, 1999) as a possible kinematic model direction were deeper (Fig. 12E and Supplemental Figs. S2–S3 [footnote 1]). The to explain the spatio-temporal evolution of these basins, which opened from

N 5000 N 5000 N P IC

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Figure 13. Comparison with nature: Recôncavo-Tucano-Jatobá basins, Brazil. (A) Cartoon illustration of the regional geological setting (top) and comparison with intermediate obliquity model E2 (α = 60°); (bottom) Blue arrow indicates the sense of rotation of the moving plate (in gray). Red arrows illustrate the regional stretching direction. Model results are rotated 35° for comparison. PL—Pernambuco Lineament; VB—Vaza Barris; MC—Mata-Catu. (B) Structural map of the Recôncavo basin (after Milani and Davison, 1988) (top) and comparison with Model E2 (bottom). Red dashed lines indicate the location of the inferred pre-existing crustal weaknesses, associated to shear zones of Pan-African age. Rose diagrams show a good match between fault distribution in nature and in the model. MC—Mata-Catu; IT—Itaporanga; IC—Itapicuru. (C) Structural map of the Jatobá basin (after Milani and Davison, 1988) and comparison with Model D (α = 45°). Example of propagating rift branch terminating abruptly against a linear crustal heterogeneity.

south to north (Fig. 13A). These conditions are comparable to experiments Davison, 1988). Such crustal anisotropies may have localized major transfer with intermediate anisotropy obliquity (Models D [α = 45°] and E [α = 60°]) zones in the Recôncavo-Tucano-Jatobá system (Fig. 13B). Two of these shear (Fig. 13A). Fault trend analysis in the southernmost Recôncavo rift basin shows zones dislocated the eastern section of the Recôncavo rift into three distinct a dominant N25°E oriented normal fault network and a secondary N30°W compartments (Milani and Davison, 1988), similar to the displacement maps trending strike-slip fault group (Milani and Davison, 1988). The final architec- for our intermediate obliquity models (Fig. 11). The overall movement sense ture of Model E2 shows a very similar fault distribution both in terms of both of reactivated shear zones and its change over time is difficult to determine in geometry and sense of motion (Fig. 13B). nature without corresponding kinematic indicators or crosscutting relationships. Differential strain measurements of experiments with α = 60° show a top- Analyzing the evolution of laboratory experiments therefore provides insights view sigmoidal strain pattern as rifts propagated across the crustal anisotropies into the complex rifting history of natural examples such as northeastern Brazil. (Fig. 7 and Supplemental Fig. S5 [footnote 1]), similar to the inferred deflecting Field observations and differences in the throw of boundary faults sup- role of pre-existing basement weaknesses in the Recôncavo rift (Milani and port a dextral strike-slip motion on the NW section of the Mata-Catu transfer

fault (Milani and Davison, 1988) (Fig. 13B). However, there are few kinematic of the MER reveal a complex tectono-magmatic evolution (e.g., Ebinger et al., constrains for its SE section. A sinistral sense of motion has been interpreted, 2000; Bonini et al., 2005; Corti, 2009). The role of two major lineaments in the based on the present-day structural pattern (Milani and Davison, 1988). Our evolution and propagation of the MER is of particular interest to this study. modeling provides additional constraints, since displacement field maps for The approximately E-W–oriented Goba-Bonga (GB) and Yerer-Tullu Wellel model E2 (α = 60°, anticlockwise) show a similar structural pattern to the (YTW) lineaments (Fig. 14A) are parallel to the general stretching direction (Abebe southeastern end of the Mata-Catu fault (Fig. 13B). The model results indicate et al., 1998; Corti et al., 2018a, 2018b). The Somalian plate has been rotating clock- constant dextral oblique- and strike-slip motion on transfer zones in the north- wise with respect to the Nubian plate since the middle Miocene (Bonini et al., 2005; ernmost sections of the experiment (Fig. 13B), as inferred for the Itapicuru, Stamps et al., 2008) (Fig. 14A). Recent work (Balestrieri et al., 2016; Boone et al., Itaporanga, and north of the Mata-Catu transfer zones (Milani and Davison, 2019) has illustrated a complex scenario of rift initiation in the MER; deformation 1988). However, the distal domain shows more distributed deformation and propagated northward in the southern MER, while the northern MER propagated a more complex evolution. At later stages in the experiment there is a local southward, into the central MER. The boundary conditions in model F (α = 90°) increase in displacement rates southwest of the southernmost crustal hetero- are therefore analogous to the tectonic setting of the northern MER (Fig. 14A). geneity (Fig. 13B). A possible explanation for this is progressive rotation of the Extensional structures developed in Southern Afar during the middle Mio- moving plate. While initially unfavorably oriented, the pre-existing weakness cene formation of the Afar triple junction but did not propagate beyond the was eventually reactivated with a dextral sense of motion (Fig. 13B). At later YTW lineament at that time (Bonini et al., 2005) (Fig. 14B). This is consistent stages, a shift in the relative orientation between the stretching direction and with the evolution of Model F, in which propagation of initial rift-parallel struc- the heterogeneity orientation may have caused a reactivation with a sinistral tures appeared to slow down when they approached the first crustal anisotropy sense of motion (Fig. 13B), which correlates well with observations in the (Fig. 14C). Extensional deformation continued to propagate southwards beyond eastern end of the Recôncavo rift (Milani and Davison, 1988). the YTW lineament and into the central MER in the late Miocene as a result The main boundary faults north of the Jatobá Graben terminate abruptly of clockwise Somalian plate rotation (Collet et al., 2000; Bonini et al., 2005) against the Ibimirim fault, which is associated with the major ENE-WSW–trend- (Fig. 14B). At this time, rift propagation was toward the SW, perpendicular to ing Pernambuco Lineament (Fig. 13C). Similar behavior is observed in Model D the general direction of extension, as observed in the Model F (Fig. 14C). It (α = 45°). Rift propagation in the central domain halted when the northernmost has been suggested that the presence of the YTW led to a minor right lateral crustal anisotropy was reached and strain was relayed onto the transfer zone offset of the rift segments (Keranen and Klemperer, 2008; Corti, 2009), which (Fig. 6), as probably occurred in the Jatobá Graben-Ibimirim fault interaction was also observed in Model F (Fig. 14C). Southwestward rift propagation zone (Milani and Davison, 1988). The North Tucano Graben has also been was again hindered during the Pliocene by the transversal GB lineament, as interpreted to cut through a basement weakness due to the latter’s unfavorable interpreted by Bonini et al. (2005) and similarly observed in Model F (Fig. 14B). orientation for reactivation (Darros de Matos, 1992). This is analogous to mod- Earlier interpretations (Bonini et al., 2005) suggested that during this period els F (α = 90°) and H (α = 75°, anticlockwise), in which the crustal anisotropies the southern MER (SMER) underwent a tectonic pause, which is consistent with did not affect the northward trend of rift propagation (Fig. 9). the evolution of our experiment. However, recent thermochronological data The inherent complexity of the northeastern Brazil rift system makes it from the SMER (Balestriei et al., 2016) indicate that rifting in the SMER activated difficult to test in a single analogue experiment. However, the intermediate prior to extensional deformation in the central MER, which indicates an even to high obliquity between the general extension direction and the pre-exist- more complex scenario. In spite of the differences observed between nature ing basement structures makes it a good reference to compare with several and experiments due to the high complexity of the Ethiopian Rift, the overall experimental results (i.e., Models D, E, F, and H). Detailed analysis of the evolution of the northern MER shows a good match with Model F, in which experiments presented here provide plausible explanations for why planar progressive rotational stretching resulted in the propagation of rift branches basement weaknesses were either ignored or followed by the evolving rifts across transversal crustal anisotropies (see Fig. 8). While the anisotropies in the Recôncavo-Tucano-Jatobá system. appear to have little influence in the orientation of the extensional structures (Fig. 12), they did play an important role stalling rift propagation, resulting in a pulsed rift history (e.g., Bosworth, 2015). In the experiment, the final stages High Obliquity (α > 60°) were characterized by the southward continuation of rift propagation across the linear crustal anisotropies (Fig. 14B). This scenario resembles earlier inter- Main Ethiopian Rift pretations of the final stages of the MER evolution during the Pleistocene, characterized by the propagation of the rift across the GB lineament and sub- The NE-SW–trending Main Ethiopian Rift (MER) is located between the Afar sequent reactivation of early Miocene structures in the SMER (Bonini et al., depression and the Kenya rift in the Eastern Branch of the East African Rift Sys- 2005) (Fig. 14B). Although the previously mentioned data from the SMER tem (Fig. 14A). The intricate network of extensional faults and volcanic history (Balestrieri et al., 2016) represents evidence against this order of segmented

Late Miocene (10-5 Ma) Pliocene (5-3 Ma) Pleistocene (3-0 Ma) Nubian Gulf Plate of Aden 10˚N Afar Afar Afar N

MER 4˚N SMER

b Somalian Plate 2˚S

300 km P

30˚E 36˚E 42˚E 6 h / 4.8 m.y. 9.3 h / 7.6 m.y. 12 h / 9.7 m.y.

c N

100 mm P DiffS [%]

Model F 0

Figure 14. Comparison with nature: Main Ethiopian Rift (MER) (East Africa). (A) Cartoon illustration of the regional geological setting (top) and comparison with high obliquity model F (α = 90°); (bottom) Blue arrow indicates the sense of rotation of the moving plate (in gray). Red arrows illustrate the regional stretching direction. Model results are mirrored and rotated ~160° for comparison. (B) Schematic evolution of the main rift segments of the MER (after Bonini et al., 2005 and Balestrieri et al., 2016). (C) Differential strain field map evolution for equivalent downscaled experimental timesteps, showing the hindering effect of the heterogeneities in rift propagation. SMER— Southern Mid-Ethiopian Rift; YTW—Yerer-Tullu Wellel lineament; GB—Goba-Bonga lineament.

rift activation events, the significant role of crustal anisotropies as obstacles heterogeneous stress fields. The use of a stereoscopic PIV system to monitor to rift propagation is still comparable. Northward propagation of the SMER, experiments provides key insights on the temporal evolution of rift develop- associated with deformation in the Kenya Rift (Bonini et al., 2005; Balestrieri ment and propagation across crustal heterogeneities. The results illustrate et al., 2016), was temporarily halted when it reached the GB lineament during how the orientation and obliquity of rheological heterogeneities in the crust the early Miocene. We therefore consider crustal anisotropies as structures that have a strong influence on fault trends and lengths, the sense of movement can cause a partial impediment to rift propagation. The development of locked on reactivated oblique-slip zones, and on the role crustal anisotropies play in zones has been identified in oceanic rifts (e.g., Courtillot, 1982; van Wijk and dislocating the crust into kinematically independent blocks. Blackman, 2005) but their origin in continental rifts remains poorly understood. The experiments show that extension-normal and low obliquity heterogeneities (α ≤ 30°) favor the formation of anomalously long normal faults, which reach their near-final length early in the evolution of rifts in comparison to ■■ CONCLUSIONS cases with a homogeneous lithosphere. Low to intermediate obliquity angles (~15° ≤ α ≤ 45°) oriented clockwise with respect to the extension direction Kinematic rotation in three-dimensional laboratory experiments simu- favor sinistral oblique-slip reactivation of linear anisotropies and subsequent lates realistic initial conditions for natural extensional settings with laterally dislocation of the crust into blocks with different displacement rates. Such

dislocation is most effective and long lived when α = 30°. Conversely, compa- Benes, V., and Scott, S., 1996, Oblique rifting in the Havre Trough and its propagation into the continental margin of New Zealand: Comparison with analogue experiments: Marine Geo- rable obliquity angles with an anticlockwise orientation results in conditions physical Researches, v. 18, p. 189–201, https://doi.org/10.1007/BF00286077. that favor reactivation of heterogeneities as dip-slip or dextral oblique-slip Boone, S.C., Balestrieri, M.‐L., Kohn, B.P., Corti, G., Gleadow, A.J.W., and Seiler, C., 2019, Tectono- normal faults. Nevertheless, the amount of oblique-slip or dip-slip deformation thermal evolution of the Broadly Rifted Zone, Ethiopian Rift: Tectonics, v. 38, p. 1070–1100, https://doi.org/10.1029/2018TC005210. in all experiments varies over time and specific interpretations must be made Bonini, M., Souriot, T., Boccaletti, M., and Brun, J.P., 1997, Successive orthogonal and oblique for each particular scenario. Overall, our results demonstrate that the lateral extension episodes in a rift zone: Laboratory experiments with application to the Ethiopian sense of motion on faults is determined by the relative orientation of crustal Rift: Tectonics, v. 16, p. 347–362, https://doi.org/10.1029/96TC03935. Bonini, M., Corti, G., Innocenti, F., Manetti, P., Mazzarini, F., Abebe, T., and Pecskay, Z., 2005, Evo- anisotropies with respect to the stretching direction. Finally, very high obliq- lution of the Main Ethiopian Rift in the frame of Afar and Kenya rifts propagation: Tectonics, uity crustal anisotropies (α > 60°) have little influence on the trends of border v. 24, https://doi.org/10.1029/2004TC001680. faults in propagating rifts. However, they may cause pulses in along-axis rift Bosworth, W., 2015, Geological evolution of the Red Sea: Historical background, review, and synthesis, in Rasul, N.M.A., and Stewart, I.C.F., eds., The Red Sea: The Formation, Morphology, propagation, which can subsequently result in variations in subsidence rates, Oceanography and Environment of a Young Ocean Basin: Heidelberg, Germany, Springer, fault throws and depocenter depths along the rift. p. 45–78, https://doi.org/10.1007/978-3-662-45201-1_3. 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