50 years of plate tectonics: then, now, and beyond

50 years of plate tectonics: then, now, and beyond

Even if we cannot attend all conferences ourselves, your EGU GD Blog Team has reporters that make sure all significant geodynamics events are covered. Today, Marie Bocher, postdoc at the Seismology and Wave Physics group of ETH Zürich, touches upon a recent symposium in Paris that covered one of the most important milestones of geodynamics.

On the 25th and 26th of June, the Parisian Collège de France was celebrating the anniversary of the plate tectonics revolution with a symposium entitled 50 years of plate tectonics: then, now and beyond. For this occasion, the organizers Eric Calais, Anny Cazenave, Claude Jaupart, Serge Lallemand, and Barbara Romanowicz had put together a very impressive list of presenters, starting with Xavier Le Pichon, Jason Morgan, and Dan McKenzie during the first morning!

The very impressive program of the 50 years plate tectonics symposium

Needless to say, it was a blast, and a great occasion to focus on the big picture and reflect on the evolution of Earth sciences within the last 50 years.

Watch it online!

But don’t panic if you missed it: all the presentations are available online now on the Collège de France website. So relax, brew yourself a cup of coffee, and enjoy the symposium from the comfort of your own home 🙂

Xavier Le Pichon
Image courtesy of Martina Ulvrova

Important panel
Image courtesy of Martina Ulvrova

Dietmar Müller
Image courtesy of Marie Bocher

Subduction through the mantle transition zone: sink or stall?

Subduction through the mantle transition zone: sink or stall?

The Geodynamics 101 series serves to showcase the diversity of research topics and methods in the geodynamics community in an understandable manner. We welcome all researchers – PhD students to professors – to introduce their area of expertise in a lighthearted, entertaining manner and touch upon some of the outstanding questions and problems related to their fields. For our latest ‘Geodynamics 101’ post, Saskia Goes, Reader at Imperial College London, UK, discusses the fate of subducting slabs at the mantle transition zone.

Saskia Goes

Subducting plates can follow quite different paths in their life times. While some sink straight through the upper into the lower mantle, others appear to stall in the mantle transition zone above 660 km depth. Geodynamicists have long puzzled about what controls these different styles of behaviour, especially because there appear to be correlations between sinking or stalling with faster or slower plate motions and mountain building or ocean basin formation, respectively. In the long run, how easily slabs sink through the transition zone controls how efficiently material and heat are circulated in the mantle.

The word subduction derives from the Latin verb subducere, which means pulled away from below, but metaphorically can mean to lose footing or remove secretly. Definitely, when Wegener first proposed continental drift, people were unaware that subduction is removing plates from the Earth’s surface. We now know this process is not quite so secret. The plates creak in earthquakes as they sink into the mantle, in some cases all the way through the mantle transition zone to about 700 km depth. Furthermore, where the subducting plate bends below the overriding plate, it creates deep-sea trenches with prominent gravity and geoid signals. This bending is a very important part of subduction dynamics, as I’ll explain below.

The seismic Wadati-Benioff zones and gravity expressions were sufficient clues of the location of the downwelling limbs of a mantle convection system to help acceptance of plate tectonics in the 1960s. However, it took another twenty odd years until seismology yielded images of cold plates sinking into the mantle, and it turned out that the plates extend beyond the seismic Wadati-Benioff zones [Van der Hilst et al., 1991; Zhou and Clayton, 1990]. These images showed that some subducting plates flatten in the mantle transition zone (e.g. below Japan and Izu-Bonin), while others continue with little to no deflection into the lower mantle (e.g., below the Northern Kuriles and Marianas) (Fig. 1). Soon after, it was realised that many of the places where the slabs are flat in the transition zone have a history of trench retreat [Van der Hilst and Seno, 1993]. Furthermore, mapping of seafloor ages revealed that flat slabs tend to form where plates older than about 50 Myr are subducted [Karato et al., 2001; King et al., 2015].

Figure 1: Variable modes of slab-transition zone interaction

Many mechanisms have been proposed for the variable slab transition-zone interaction. We recently reviewed the geodynamic and observational literature and combined these insights with those from our own set of mechanical and thermo-mechanical subduction models [Goes et al., 2017]. This effort shows that not one single mechanism, but an interplay of several mechanisms is the likely cause of the observed variable subduction behaviour.

It has long been realised that viscosity increases with depth into the mantle, quite possibly including jumps at the major phase transitions in the mantle transition zone. The ringwoodite-postspinel transition that is responsible for the global 660 km seismic discontinuity, usually taken as the base of the upper mantle, is an endothermic transition under most of the conditions prevailing in the mantle today. This means that the transition will take place at a higher pressure and thus depth in the subducting plate than the surrounding mantle, rendering the plate locally buoyant with respect to the mantle. Both these factors hamper the descent of the subducting plate through the transition zone. However, a viscosity increase within acceptable bounds (as derived from geoid and postglacial rebound modelling) can slow sinking, but does not lead to stalling material. By contrast, the phase transition can lead to stalling, as well as an alternation of periods of accumulation of material in the transition zone and periods where this material flushes rapidly into the lower mantle, at least in convection models without strong plates. But does this work with strong plates?

Making dynamic models of subduction with strong plates is challenging because the models need to capture strong strength gradients between the core of the plate and the underlying mantle, allow for some form of plate yielding, maintain a weak zone between the two plates and adequately represent the effect of plate bending (a free-surface effect). Most models prescribe at least part of the system by imposing velocities and/or plate geometries. This however needs to be done with great care and consideration for what forcing such imposed conditions imply.

“Pulled away from below” is a good description of the dynamics of subduction. Subduction is primarily driven by slab pull, the gravitational force on the dense subducting plate [Forsyth and Uyeda, 1975]. And to “lose footing” reminds us that gravity is the main driving force. Gravity tries to pull the plate straight down (Fig. 2), so the easiest way for a plate to subduct is to fall into the mantle, a process that leads to trench retreat [Garfunkel et al., 1986; Kincaid and Olson, 1987]. Besides letting the plate follow the path of gravity, subduction by trench retreat has the other advantage that the plate does not need to bend too much. Bending a high-strength plate takes significant energy. Some studies have shown that if plates are assigned laboratory-based rheologies, such bending can easily take up all of the gravitational potential energy of the subducting plate [Conrad and Hager, 1999], so if plates are to sink into the mantle, they have to do this by minimising the amount of energy used for bending into the trench. As a consequence, strong and dense plates prefer to subduct at smaller dip angles while weaker and lighter plates can be bent to subduct more vertically [Capitanio et al., 2007].

Figure 2: If subduction occurs freely, i.e., driven by the pull of gravity on the dense slab with sinking resisted by the viscous mantle, it is usually energetically most favourable to subduct by trench retreat.

The angle at which plates subduct strongly affects how they subsequently interact with viscosity or phase interfaces (Fig. 3). Steeply dipping plates will buckle and thicken when they encounter resistance to sinking. This deformation facilitates further sinking, as a bigger mass. But plates that reach the interface at a lower dip may be deflected. Such deflected plates have a harder time sinking onwards, both because the high viscosity resistance is now distributed over a wider section of the plate and due to the spread-out additional buoyancy from the depressed endothermic phase boundary.

Figure 3: The subduction angle largely determines how the slab interacts with viscosity and phase changes.

So, variable plate density and strength can lead to variable behaviour of subduction in the transition zone. And we know plates have variable density and strength. Older plates are denser and if strength is thermally controlled, as most lab experiments predict, also stronger than younger plates. This implies that older plates can drive trench retreat more easily than young plates. And indeed this matches observations that significant trench retreat has only taken places where old plates subduct. Furthermore, significant trench retreat will facilitate plate flattening in the transition zone, consistent with the observation that flat plates tends to underlie regions with a history of trench retreat (even if that does not always mean trench motions are high at the present day). This mechanism can also explain why flat slabs tend to be associated with old plate subduction.

So what about the role of other proposed mechanisms? Our models with strong slabs show that only when slabs encounter both an increase in viscosity (which forces the slabs to deform or flatten) and an endothermic phase transition (which can lead to stalling of material in the transition zone) do we find the different modes of slab dynamics. Neither a viscosity increase alone, nor an endothermic phase transition alone leads to mixed slab dynamics.

Other factors likely contribute to the regional variability. In the cold cores of the slabs, some phases may persist metastably, thus delaying the transformations to higher density phases to a larger depth. Metastability will be more pervasive in colder old plates thus making older plates more buoyant and hence resistant to sinking than young ones. In combination with trench retreat facilitated by a strong slab at the trench, this can further encourage slab flattening [Agrusta et al., 2014; King et al., 2015]. Phase transformations may also lead to slab weakening in the transition zone because they can cause grain size reduction. Such weakening can aid slab deflection [Čížková et al., 2002; Karato et al., 2001]. However, several studies have shown that transition zone slab strength is less important than slab strength at the trench, which governs how a slab starts sinking through the transition zone.

The Earth is clearly more complex than the models discussed. For example, present-day plate dip angles display various trends with plate age at the trench. Lateral variations in plate strength and buoyancy can complicate subduction behaviour. Furthermore, forces on the upper plate and large-scale mantle flow may also impede or assist trench motions and may thus affect or trigger changes in how slabs interact with the transition zone [Agrusta et al., 2017]. All these factors remain to be fully investigated. However, the first order trends of subduction-transition zone interaction can be understood as a consequence of plates of various ages interacting with a viscosity increase and endothermic phase change.

 Agrusta, R., J. van Hunen, and S. Goes (2014), The effect of metastable pyroxene on the slab dynamics, Geophys. Res. Lett., 41, 8800-8808.
 Agrusta, R., S. Goes, and J. van Hunen (2017), Subducting-slab transition-zone interaction: stagnation, penetration and mode switches, Earth Planet. Sci. Let., 464, 10-23.
 Capitanio, F. A., G. Morra, and S. Goes (2007), Dynamic models of downgoing plate buoyancy driven subduction: subduction motions and energy dissipation, Earth Planet. Sci. Lett., 262, 284-297.
 Čížková, H., J. van Hunen, A. P. van der Berg, and N. J. Vlaar (2002), The influence of rheological weakening and yield stress on the interaction of slabs with the 670 km discontinuity, Earth Plan. Sci. Let., 199(3-4), 447-457.
 Conrad, C. P., and B. H. Hager (1999), Effects of plate bending and fault strength at subduction zones on plate dynamics, J. Geophys. Res., 104(B8), 17551-17571.
 Forsyth, D. W., and S. Uyeda (1975), On the relative importance of driving forces of plate motion. , Geophys. J. R. Astron. Soc. , 43, 163-200.
 Garfunkel, Z., C. A. Anderson, and G. Schubert (1986), Mantle circulation and the lateral migration of subducted slab, J. Geophys. Res., 91(B7), 7205-7223.
 Goes, S., R. Agrusta, J. van Hunen, and F. Garel (2017), Subduction-transition zone interaction: a review, Geosphere, 13(3. Subduction Top to Bottom 2), 1-21.
 Karato, S. I., M. R. Riedel, and D. A. Yuen (2001), Rheological structure and deformation of subducted slabs in the mantle transition zone: implications for mantle circulation and deep earthquakes, Phys. Earth Plan. Int., 127, 83-108.
 Kincaid, C., and P. Olson (1987), An experimental study of subduction and slab migration, J. Geophys. Res., 92(B13), 13,832-813,840.
 King, S. D., D. J. Frost, and D. C. Rubie (2015), Why cold slabs stagnate in the transition zone, Geology, 43, 231-234.
 Van der Hilst, R. D., and T. Seno (1993), Effects of relative plate motion on the deep structure and penetration depth of slabs below the Izu-Bonin and Mariana island arcs, Earth Plan. Sci. Let., 120, 395-407.
 Van der Hilst, R. D., E. R. Engdahl, W. Spakman, and G. Nolet (1991), Tomographic imaging of subducted lithosphere below northwest Pacific island arcs, Nature, 353, 37-43.
 Zhou, H.-w., and R. W. Clayton (1990), P and S Wave Travel Time Inversions for Subducting Slab Under the Island Arcs of the Northwest Pacific, J. Geophys. Res., 95(B5), 6829-6851.

Remarkable Regions – The India-Asia collision zone

Remarkable Regions – The India-Asia collision zone

Every 8 weeks we turn our attention to a Remarkable Region that deserves a spot in the scientific limelight. This week we zoom in on a particular part of the eastern Tethys as Adina Pusok discusses the India-Asia collision zone. She is a postdoctoral researcher at the Institute of Geophysics and Planetary Physics, Scripps Institution of Oceanography, UCSD, US.

Without doubt, one of the most striking features of plate tectonics and lithospheric deformation on Earth is the India-Asia collision zone, largely comprised of the Himalayan and Karakoram mountain belts and the Tibetan plateau. What makes this collision zone so remarkable? For one, Tibet is the largest, highest and flattest plateau on Earth with an average elevation exceeding 5 km, and it includes over 80% of the world’s land surface higher than 4 km. Then, the bordering Himalayas and the Karakoram Mountains include the only peaks on Earth reaching more than 8 km above sea level.

It makes one wonder, how can such a mountain belt and high plateau form? Most of the major mountain belts and orogenic plateaus on Earth are found within the overlying plate of subduction and/or collision zones (e.g. the Alps, the Andes, the American Cordilleras etc.). When an ocean closes and two continental plates meet at a destructive (subduction) boundary, the continents themselves collide. Such collisions result in intense deformation at the edges of the colliding plates. Neither continent can be subducted into the mantle due to the buoyancy of continental crust, so the forces that drive the plate movement prior to collision are brought to act directly on the continental lithosphere itself. At this stage, further convergence of the plates must be taken up by deforming one or both of the plates of continental lithosphere over hundreds of kilometres [Figure 1]. Mountain belts can form under these circumstances.

Figure 1 Global map of surface velocities and the second invariant of strain rate (from Moresi [2015]). The surface velocities show the location and extent of plates, and the strain rate map highlights the fact that most of the deformation is concentrated at plate boundaries (high strain rates), while the continental interiors have little or no deformation (low strain rates). In some places, deformation occurs over broader regions, especially following mountain belts. These boundaries are called diffuse plate boundaries. The white rectangle roughly indicates the extent of the India-Asia collision zone.

The Himalayas and the Tibetan plateau are no different. Following the closure of the Tethys ocean (see earlier blog post), the Indian continent collided with Eurasia around 50 million years ago (e.g., Patriat and Achache [1984]), thus giving rise to this anomalously high region. This tectonic boundary is complex and changes character along its length. The Tibetan plateau is a collage of continental blocks (terranes) that were added successively to the Eurasian plate during the Paleozoic and Mesozoic [Figure 2]. The boundaries between these terranes are marked by scattered occurrences of ophiolitic material, which are rocks characteristic of oceanic lithosphere. The Himalayas represent the traditional accretionary wedge formed by folding and thrusting of sediments scraped off the subducting slab.

Figure 2 Simplified tectonic map of Tibet and surrounding region showing approximate boundaries of the major terranes, suture zones, and strike-slip faults (from Ninomiya and Bihong [2016]). Blocks and terranes: ALT-EKL-QL: Altyn Tagh–East Kunlun–Qilian terrane; BS: Baoshan terrane; HM: Himalayan terrane; IC: Indo-China block; KA: Kohistan Arc terrane; LA: Ladakh arc terrane; LC: Lincang–Sukhothai–Chanthaburi Arc terrane; LST: Lhasa terrane; NCB: North China Block; NQT: North Qiangtang terrane; QT: Qiangtang terrane; SP: South Pamir terrane; SPGZ: Songpan-Ganze terrane; SQT: South Qiangtang terrane; TC: Tengchong terrane; TSH: Tianshuihai terrane; WB: West Burma terrane; WKL: West Kunlun terrane. Suture zones: BNS: Bangong-Nujiang Suture; EKLS: East Kunlun Suture; ITS: Indus-Tsangbo Suture; JSS: Jinsha Suture; LSS: Longmu Tso–Shuanghu–Menglian–Inthanon Suture; WKLS: West Kunlun Suture. Basins: QB: Qaidam Basin; KB: Kumkol Basin. Faults: ALT: Altyn Tagh Thrust; ALTF: Altyn Tagh Fault; KKF: Karakorum Fault; LMST: Longmen Shan Thrust; MFT: Main Frontal Thrust; NQLT: North Qilian Thrust; RRF: Red River Fault; SGF: Sagaing Fault; XXF: Xianshui River–Xiaojiang Fault.

Interestingly, the India-Asia collision orogen is not just the youngest and most spectacular active continent collision belt, it is also the most studied research area on Earth. Studies on this region span a wide range of topics and methods for over more than 100 years. I am not sure if it is the fascination with the highest mountain on Earth (Mt. Everest was actually climbed for the first time as late as 1953 by Tenzing Norgay and Edmund Hillary), similar to our fascination for exploring the Moon, Mars and the other planets in our Solar System nowadays, or the hope that studying the youngest orogeny will help us decipher the older ones (soon to realize different mountain belts evolve differently).

To understand the magnitude of the work done in the past 100 years, a simple search of the keywords “India Asia collision” on Google Scholar yielded ~90k results, and a more focused geosciences search on Web of Science (where I filtered the results to those from geophysics, geochemistry, geology, geosciences multidisciplinary only) yielded >1600 results for the same keywords (other keywords: “Himalaya” > 5600 results, “Tibet” > 6500 results, “India Asia” > 2200 results). These numbers can be intimidating to a new student taking on the topic, but it is a topic worth studying and I’ll explain why below.

From a general perspective, it is important to study the India-Asia collision zone due to the interaction between tectonics and climate and the formation of the Indian monsoon [Molnar et al., 1993], but also because it is a highly populated area (>200 million people in the Hindu Kush Himalaya region) regularly shaken by natural phenomena, such as earthquakes, floods or landslides. For example, the last large earthquake in Nepal, the Gorkha earthquake (Mw 7.8) in April 2015 caused more than 9000 deaths.

From a geophysics point of view, understanding mountain-building processes and the driving forces of plate tectonics has been one of the long-term goals of solid Earth sciences community. The India-Asia collision zone is one of the best examples in which subduction, continental collision and mountain building can be studied in a global plate tectonics perspective. Prior to plate tectonics theory, Argand [1924] and Holmes [1965] thought that the Himalayan Mountains and Tibetan Plateau had been raised due to the northern edge of the Indian craton underthrusting the entire region, causing shortening and thickening of the crust to ∼80 km. This perspective remains widely accepted, but recent ideas suggest that other processes are equally important (more below).

Today, the challenge lies in refining our understanding of the dynamics of India-Asia collision by elucidating the connections between the wealth of observations available and the underlying processes occurring at depth. Decades of study have produced data sets across various disciplines, including: active tectonics, Cenozoic geology, seismicity, global positioning system (GPS) measurements, seismic profiles, tomography, gravity anomalies, mantle-crustal anisotropy, paleomagnetism, geochemistry or magnetotelluric studies. Of these, the GPS data stands out as it clearly shows the distributed deformation across the entire collision zone and suggests that this is a highly dynamic area [Figure 3].

Figure 3 Horizontal GPS velocities of crustal motion around the Tibetan Plateau relative to stable Eurasia from Liang et al. [2013].

Collectively, all these observation data sets stand as a different piece in the puzzle of the India-Asia collision. However, the same data sets can support a number of competing and sometimes mutually exclusive mechanisms for the uplift of the Tibetan Plateau. For example, the mantle lithosphere beneath Tibet has been proposed to be cold, hot, thickened by shortening, or thinned by viscous instability. Other controversies include the degree of mechanical coupling between the crust and deeper lithosphere and the nature of large-scale deformation. It is no surprise then, that several hypotheses emerged over time trying to explain the anomalous rise of the Himalayas and Tibetan Plateau [Figure 4]:

  1. Figure 4 Schematic cartoons of tectonic models proposed to explain the thickening and uplift of the Himalayas and the Tibetan Plateau. (Source: personal institutional web page of A. Ozacar).

    Wholescale underthrusting of the Indian plate below the Asian continent [e.g. Argand, 1924].

  2. The thin-sheet model or distributed homogeneous shortening [e.g. England and McKenzie, 1982].
  3. Homogeneous thickening of a weak, hot Asian crust, involving a large amount of magmatism [e.g. Dewey and Burke, 1973].
  4. Slip-line field model to account for the brittle deformation in and around the Tibetan Plateau and to explain extrusion of SE Tibet away from Indian continent [e.g. Molnar and Tapponnier, 1975]. The same group also proposes a time-dependent model for the growth of Tibetan plateau [e.g. Tapponnier et al., 2001], in which successive intracontinental subduction zones maintain the stepwise growth and rise of the plateau.
  5. Lower crustal flow models for the exhumation of the Himalayan units and lateral spreading of the Tibetan plateau [e.g. Royden et al., 1997, Beaumont et al., 2001].
  6. Delamination or convective removal of the lithospheric mantle that induced isostatic movement, lifting the Tibetan Plateau [e.g. Molnar, 1988].


These models were applied either to the Tibetan Plateau or the Himalayan mountain belt and were able to explain the formation of specific tectonic and geological features. However, there is no conclusive answer on which of the hypotheses works best for the entire orogen, and instead, more questions arise:

  • Which forces are at work during continental collision and mountain building?
  • What is the deformation history and evolution of this plate boundary?
  • How was the subduction accommodated in the Neo-Tethys?
  • How does subduction evolve during continental collision?
  • What drives the present-day fast convergence (~4-5 cm/yr) between India and Eurasia?
  • Which forces propagated India northwards between 70-50 million years at anomalously high speeds (up to 16 – 20cm/yr)?
  • How can you form such large elevations over such extended areas?
  • What is the effect of surface processes on uplift?
  • What is the structure at depth beneath the Himalayas and Tibetan plateau?
  • How do the Indian and Eurasia plates deform during collision?
  • How is the deformation accommodated during continental collision?
  • How do mountain belts form and why not all mountain belts look the same?
  • How did the crust beneath Himalaya and Tibet reach double-crustal thickness (normal continental crust is 35-40 km thick, whereas the crust beneath the Himalaya and Tibet is 70-100 km thick)?
  • Which mechanisms help sustain the high topographic amplitudes?
  • Why should an area as broad as the Tibetan Plateau be uplifted so high compared to other mountain belts following collision?
  • Did the Tibetan Plateau and Himalayan mountain belt rise continuously or diachronously?
  • Which the proposed models [Figure 4] can be applied, and where?
  • How do lithospheric heterogeneities and rheology affect the deformation pattern?
  • What is the degree of mechanical coupling between the crust and deeper lithosphere? Is it the “jelly sandwich” model (e.g., Burov and Watts [2006]) or the “creme-brulee” model (e.g., Jackson [2002], see earlier blog post)?
  • Why do the Himalayas have a convex curvature?
  • What about the high deformation of the prominent Himalayan syntaxes (the inflection points of the Himalayan belt): Nanga Parbat in the west and Namche Barwa in the east?
  • What is the effect of the India-Asia collision on climate? Do the Himalayas affect the Indian monsoon or is it the other way around? A chicken-and-egg question.

Seriously, can I even stop asking questions? The question that fascinated me the most during my graduate studies was “Why is the Himalayan-Tibet region so high and broad compared to other mountain belts?”. If we tune our models to Earth parameters, can we build such large elevations in computer simulations? Which factors and forces are at play? Using 3-D numerical models to address this question [Pusok and Kaus, 2015], we were able to obtain distinct topographic modes (different types of mountain belts) [Figure 5] and to show that building topography is an interplay between providing the energy to the system and the ability of that system to store it over longer periods of time. We also suggest that the reason why Himalaya-Tibet is different from the Alps, for example, is because the shape and elevation of mountain ranges can vary depending on the boundary conditions (plate driving forces that control convergence velocity and lithospheric heterogeneities such as the Tarim Basin) and internal factors (rheology), but also on the evolution stage they are in.

To sum up, it is clear that many of the above questions remain unanswered. But I think this is good news, meaning that in the future, exciting new results will shape our understanding of this remarkable region.

Figure 5 3-D Simulation results showing different modes of surface expressions in continental collision models. Modified from Pusok and Kaus [2015].

Argand, E. (1924). La tectonique de l’Asie. Proc. 13th Int. Geol. Cong., 7:171–372.

Beaumont, C., Jamieson, R. A., Nguyen, M. H., and Lee, B. (2001). Himalayan Tectonics Explained by Extrusion of a Low-Viscosity Crustal Channel Coupled to Focused Surface Denudation. Nature, 414:738–742.

Burov, E. B. and Watts, A. B. (2006). The long-term strength of continental lithosphere: “jelly sandwich” or “crème brûlée”? GSA Today, 16(1):4.

Dewey, J. F. and Burke, K. (1973). Tibetan, Variscan, and Precambrian Basement Reactivation: Products of Continental Collision. The Journal of Geology, 81(6):683–692.

England, P. and McKenzie, D. (1982). A Thin Viscous Sheet Model for Continental Deformation. Geophys. J. R. astr. Soc., 70:295–321.

Holmes, A. (1965). Principles of Physical Geology. The Ronald Press Company, New York, second edition.

Jackson, J. (2002). Strength of the Continental Lithosphere: Time to Abandon the Jelly Sandwich? GSA Today, 4–9.

Liang, S., Gan, W., Shen, C., Xiao, G., Liu, J., Chen, W., Ding, X., and Zhou, D. (2013). Three-dimensional velocity field of present-day crustal motion of the Tibetan Plateau derived from GPS measurements. Journal of Geophysical Research: Solid Earth, 118:1–11.

Molnar, P. and Tapponnier, P. (1975). Cenozoic Tectonics of Asia: Effects of a Continental Collision. Science, 189:419–426.

Molnar, P. (1988). A Review of Geophysical Constraints on the Deep Structure of the Tibetan Plateau, the Himalaya and the Karakoram, and their Tectonic Implications. Philosophical Transactions of the Royal Society A: Mathematical, Physical and Engineering Sciences, 326(1589):33–88.

Molnar, P., England, P., and Martinod, J. (1993). Mantle Dynamics, Uplift of the Tibetan Plateau, and the Indian Monsoon. Reviews of Geophysics, 31:357–396.

Moresi, L. (2015). Computational Plate Tectonics and the Geological Record in the Continents. SIAM News, 48:1–6.

Ninomiya, Y. and Bihong Fu, B. (2016). Regional Lithological Mapping Using ASTER-TIR Data: Case Study for the Tibetan Plateau and the Surrounding Area. Geosciences 2016, 6(3), 39; doi:10.3390/geosciences6030039.

Patriat, P. and Achache, J. (1984). India-Eurasia collision chronology has implications for crustal shortening and driving mechanism of plates. Nature, 311:615–621.

Pusok, A. E. and Kaus, B. J. P. (2015). Development of topography in 3-D continental-collision models. Geochemistry, Geophysics, Geosystems, 16(5):1378–1400.

Royden, L. H., Burch el, B. C., King, R., Wang, E., Chen, Z., Shen, F., and Liu, Y. (1997). Surface Deformation and Lower Crustal Flow in Eastern Tibet. Science, 276(5313):788–790.

Tapponnier, P., Zhiqin, X., Roger, F., Meyer, B., Arnaud, N., Wittlinger, G., and Jingsui, Y. (2001). Oblique Stepwise Rise and Growth of the Tibet Plateau. Science, 294(5547):1671–1677.


Alaska: a gold rush of along strike variations

Alaska:  a gold rush of along strike variations

Every 8 weeks we turn our attention to a Remarkable Region that deserves a spot in the scientific limelight. After exploring the Mediterranean and the ancient Tethys realm, we now move further north and across the Pacific to the Aleutian-Alaska subduction zone. This post was contributed by Kirstie Haynie who is a PhD candidate at the department of geology at the University at Buffalo, State University of New York, in the United States of America.

Given that Alaska is a remarkable region, I decided to walk up to strangers and ask them what comes to mind when they hear the word “Alaska”. Indeed I received some confusing looks and laughs, but everyone I asked had something to say. Some people alluded to popular TV shows set in Alaska, such as Gold Rush, Bush People, and Alaska: the Last Frontier, while others spoke about the cold weather, dog mushing, Eskimos, fishing and hunting, and the Trans-Alaska pipeline. A few of the answers I received referenced the beauty and wilderness of the large snow capped mountains, glaciers, and the Northern Lights (Aurora Borealis): all emblematic of the largest state in America. But to me, Alaska is more than just a pretty landscape and a place to fish. It is a region riddled with geologic mysteries and rich in along strike variations.

The Aleutian-Alaska subducton zone marks a North American-Pacific plate boundary where subduction varies greatly along strike (Figure 1). At the western end of the subduction zone, the Aleutian volcanic islands are the result of oceanic-oceanic subduction while in the eastern part of the subduction zone there is oceanic-continental collision where the Pacific plate descends beneath the North American plate. The age of the subducting sea floor increases laterally from around 30 Ma in the eastern subduction corner to 80 Ma at the end of the Aleutian volcanic arc (Müller et al., 2008). Slab dip changes drastically from 50° to 60° in the west and central Aleutians to flat slab subduction under south-central Alaska (Ratchkovski and Hansen, 2002a; Lallemand et al., 2005; Jadamec and Billen, 2010). This leads to a variation in the slab pull force, which is a main driving force of subduction caused by the weight of dense slabs sinking into the mantle (Morra et al., 2006).

Figure 1: Tectonic map of Alaska modified from Haynie and Jadamec (2017). Topography/bathymetry is from Smith and Sandwell (1997) and Seafloor (SF) ages are from Müller et al. (2008). Blue lines are the slab contours of Jadamec and Billen (2010) in 40 km intervals; the thick black line is the plate boundary from Bird (2003); and the thinner black lines are faults from Plafker et al. (1994a). The location of Denali is marked by the orange hexagon. Holocene volcanoes are given by the pink triangles (Alaska Volcano Observatory). The purple polygon is the outline of the Yakutat oceanic plateau (Haynie and Jadamec, 2017). WB – Wrangell block fore-arc sliver; JdFR – Juan de Fuca Ridge.

There is also a distinct change in margin curvature from convex in the west to concave in the east. At the end of the eastern bend, the Alaska part of the subduction zone is truncated by a large transform boundary, the Fairweather-Queen Charolette fault, which gives rise to a corner-shaped subduction-transform plate boundary (Jadamec et al., 2013; Haynie and Jadamec, 2017). Here, convergence is oblique with an average velocity of 5.2 cm/year northwest (DeMets and Dixon, 1999). Seismic studies (Page et al., 1989; Ferris et al., 2003; Eberhart-Phillips et al., 2006; Fuis et al., 2008) show that thicker than normal oceanic crust lies off-shore in the subduction corner. This thick oceanic material has been identified as the Yakutat oceanic plateau (Plafker et al., 1994a; Brocher et al., 1994; Bruns, 1983; Worthington et al., 2008; Christeson et al., 2010; Worthington et al., 2012). Even though oceanic plateaus tend to resist subduction (Cloos, 1993; Kerr , 2003), the Yakutat plateau is currently subducting beneath the Central Alaska Range to depths of 150 km (Ferris et al., 2003; Eberhart-Phillips et al., 2006; Wang and Tape, 2014). It is also colliding into south-east Alaska (Mazzotti and Hyndman, 2002; Elliott et al., 2013; Marechal et al., 2015) where the largest coastal mountain range on Earth, the Saint Elias Mountains, are located (Enkelmann et al., 2015).

With regards to surface deformation, in addition to Denali (the tallest mountain in North America), other notable along strike variations reside within the broad deformation zone of south-central Alaska. For example, a normal volcanic arc occurs over the Aleutian part of the subduction zone and above the Alaska Peninsula. However, above the flat slab there is a gap in volcanism followed by the presence of the enigmatic Wrangell volcanoes (Rondenay et al., 2010; Jadamec and Billen, 2012; Martin-Short et al., 2016; Chuang et al., 2017). These volcanoes are marked by a range of morphologies as well as adakitic geochemical signatures (Richter et al., 1990; Preece and Hart , 2004), which have a petrogenesis that may be attributed to slab melting (Defant and Drummond , 1990; Peacock et al., 1994; Castillo, 2006, 2012; Ribeiro et al., 2016). Analogue (Schellart , 2004; Strak and Schellart , 2014) and 3D numerical models (Stegman et al., 2006; Piromallo et al., 2006; Jadamec and Billen, 2010, 2012) predict that toroidal flow can produce upwellings around the edge of a slab that may have implications for melting of the slab and the formation of adakites. However, the formation of the Wrangell volcanoes is still debated.

Also located above the subducting plateau and flat slab is the Wrangell block fore-arc sliver, which exhibits northwest motion and counterclockwise rotation (Cross and Freymueller, 2008; Freymueller et al., 2008; Bemis et al., 2015; Waldien et al., 2015; Jadamec et al., 2013; Haynie and Jadamec, 2017). This sliver is bounded in the north by the arcuate shaped Denali fault, which illustrates a lateral change in slip rates that increases towards the center of the fault (Haynie and Jadamec, 2017; Haeussler et al., 2017). 3D high-resolution geodynamic models show that the flat slab drives motion of the Wrangell block fore-arc sliver (Jadamec et al., 2013; Haynie and Jadamec, 2017) and contributes to fault parallel motion along the eastern Denali fault and convergence along the apex of the fault (Haynie and Jadamec, 2017) (Figure 2). However, when model predictions of the Wrangell block motion and the difference in Denali fault parallel motion are compared with observations, model predictions are lower, suggesting that the flat slab alone is not sufficient enough to explain the broad deformation zone of Alaska (Haynie and Jadamec, 2017). Thus, it is thought that the neotectonics of south-central Alaska are predominantly driven by the subduction-collision of the buoyant Yakutat oceanic plateau (Bird , 1988; Plafker et al., 1994b; Fitzgerald et al., 1995; Ratchkovski and Hansen, 2002b; Bemis and Wallace, 2007; Chapman et al., 2008; Haeussler , 2008; Jadamec et al., 2013; Lease et al., 2016; Haynie and Jadamec, 2017). 4D numerical modelling of this process is currently underway.

Figure 2: Top: map of south-central Alaska (zoomed in from Figure 1) with model predicted velocities (blue arrows) from Haynie and Jadamec (2017) plotted on top. Bottom: percent of slab contribution from Haynie and Jadamec (2017) models to observed Denali fault slip rates (modified from Haynie and Jadamec (2017)). Results from Haynie and Jadamec (2017) show that the slab drives northwest and counter-clockwise motion of the Wrangell block fore-arc sliver and contributes to an average of 20-28% of motion along the Denali fault. The flat slab exerts the largest contribution to motion along the eastern segment of the fault, where surface motion parallels the fault, and also along the central segment of the fault, where the slab is driving the Wrangell block into the North American backstop and subducting obliquely to the fault.


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