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Remarkable Regions

Remarkable Regions – The Réunion Hotspot

Remarkable Regions – The Réunion Hotspot
Eva Bredow at Réunion caldera.

Eva Bredow in front of the caldera at Réunion Island. Credit: Simon Stähler.

This week we again turn our attention to a Remarkable Region that deserves a spot in the scientific limelight. Postdoctoral researcher Eva Bredow of Kiel University shares with us her long history with Réunion Island.

At first glance, Réunion is a relatively small tropical island, located between Madagascar and Mauritius, and from my personal experience, most Germans have never even heard of it. To be fair, it is much better known in France, because Réunion is officially a French overseas department, meaning that the eleven-hour flight from Paris is technically a domestic flight and that you can pay there with Euros (and I bet you did not know that a millimetre-sized outline of the island appears on every Euro banknote!). Besides, Réunion hosts one of the most active volcanoes in the world with one eruption per year on average. However, it rarely hits the headlines because the inhabitants live far enough away not to be overly threatened. And yet, for people interested in geodynamics, the name Réunion might actually have a familiar sound, since it regularly appears in hotspot catalogues and hotspot reference frames – a sure indication that there is more to discover.

For me, Réunion has been a very special place ever since I was a high school student lucky enough to visit the island in order to learn French. And who would have thought back then that hiking in this surreal volcanic landscape would be one of the first steps towards my decision to study geophysics? And what were the odds to stumble upon a PhD project years later, centred around the Réunion hotspot? Well, that is exactly what happened and in this article, it is my pleasure to give you at least a brief overview of why Réunion deserves to be called a remarkable spot indeed and how numerical modelling can help us to explore its geodynamic history.

NW Indian Ocean crustal thickness map.

Crustal thickness map of the north-western Indian Ocean with the entire hotspot track from Réunion Island to the Deccan Traps in India. Figure from Torsvik et al. (2013).

A deep root

The hypothesis that Réunion is an intraplate hotspot possibly fed by a hot, buoyant upwelling rooted deep in the mantle was already put forward by Jason Morgan (1971, 1972) in his famous papers outlining the classical mantle plume hypothesis. And as it happens, the Réunion plume has left a number of traces that fit the plume hypothesis extremely well and make it a kind of prototype for a deep plume and its surface manifestations. A brief look at a topographic map of the north-western Indian Ocean reveals not only the currently active hotspot at Réunion and the slightly older island of Mauritius, but also a clearly continuous (and age-progressive) hotspot track on the African and Indian plates, only split due to subsequent seafloor-spreading.

According to numerous laboratory and numerical studies that describe the mushroom-like geometry of a plume, the hotspot track is considered to be caused by the long-lived plume tail, whereas the voluminous plume head is supposed to create a huge flood basalt province in a relatively short geological time (Richards et al., 1989). In the case of the Réunion plume, the hotspot track starts at the Deccan Traps, a gigantic continental Large Igneous Province (LIP) in India. The LIP was created around 65 million years ago and the environmental changes triggered by the volcanic activities might have led to the extinction of the dinosaurs (an alternative theory to the Chicxulub impact in Mexico; Courtillot and Renne, 2003).

Further indications for a deep plume beneath Réunion include the broad topographic hotspot swell around the island, a geochemical signature of the volcanic rocks that clearly deviates from mid-ocean ridge basalts, and the present-day hotspot location above the plume generation zone at the margin of the African Large Low Shear Velocity Province (LLSVP).

Plume-ridge interaction

A more puzzling observation is the geochemical anomaly at the closest segments of the Central Indian Ridge, about 1000 km away from Réunion that implies a long-distance plume-ridge interaction. Already Morgan (1978) suggested that a sublithospheric flow channel connecting the upwelling plume and the ridge is responsible for the creation of the Rodrigues Ridge, a rather eye-catching feature not at all parallel to the hotspot track or recent plate motions.

And there is one more noteworthy hypothesis associated with Réunion, based on extremely old zircons found at Mauritius; it postulates that the hotspot track has (coincidentally) been created on top of a Precambrian microcontinent (Ashwal et al., 2017).

The RHUM-RUM experiment (completely alcohol-free…)

Concerning the (present-day) state of the Réunion plume at greater depths, seismic tomography is the most promising tool to answer the question if it is indeed fed by a deep plume or not. But given that the island is rather remotely located and a classical plume tail is expected to be quite narrow, there are plenty of technical obstacles, and it was not until 2006 that Montelli published the first seismic image of a continuous plume conduit reaching into the deep mantle. More recent global tomography models also image the Réunion plume as a clearly resolved, vertically continuous conduit at depths between 1,000 and 2,800 km (French and Romanowicz, 2015).

In 2012-2013, the French-German RHUM-RUM project (Réunion Hotspot and Upper Mantle – Réunions Unterer Mantel) aimed at an even higher resolved image of the plume. Therefore, 57 German and French ocean-bottom seismometers were deployed at the seafloor around Réunion for about a year (Stähler et al., 2016) – still the largest seismological experiment to image a deep oceanic mantle plume so far.

 

RHUM-RUM seismic stations

All seismic stations related to the RHUM-RUM project, with the 57 ocean-bottom seismometer stations shown in red. More information on the project can be found here.

With all that in mind, and as part of the RHUM-RUM project, I set up a regional numerical model with some colleagues from the GFZ Potsdam in order to assemble Réunion’s entire dynamic history. We used time-dependent plate reconstructions and large-scale mantle flow as velocity boundary conditions as well as a laterally varying lithosphere thickness in order to specifically simulate the Réunion plume (for details, see Bredow et al., 2017). In short: altogether, we were able to reproduce a crustal thickness pattern that at first order fits the observed hotspot track (although the method is not suited to reproduce a continental LIP such as the Deccan Traps). Moreover, the interaction between the plume and the Central Indian Ridge explained both the genesis of the Rodrigues Ridge and the gap in crustal thickness between the Maldives and Chagos – both features that have not been dynamically modelled before.

After our models were published, the active long-distance plume-ridge interaction beneath the Rodrigues Ridge was additionally confirmed by seismological studies in the RHUM-RUM project: first in a three-dimensional anisotropic S-wave velocity model comprising the uppermost 300 km (Mazzullo et al., 2017), and second by SKS splitting measurements (Scholz et al., 2018). Overall, these interdisciplinary studies confirmed Morgan’s long-standing hypothesis – more than 30 years after its original publication.

 

Cross section geodynamic plume model of Bredow et al. 2017.

Cross section of the geodynamic plume model, showing the long-distance plume-ridge interaction as predicted by Morgan (1978). Figure after Bredow et al. (2017).

Surface wave tomography showing the Reunion plume.

Cross section of the surface wave tomography model, showing the low velocity signature of the plume rising toward the base of the lithosphere underneath Réunion and the sublithospheric flow toward the Central Indian Ridge (CIR). Figure after Mazzullo et al. (2017).

The whole-mantle P- and S-wave tomography models from the RHUM-RUM project have yet to be published, but the (almost final) results presented at this year’s EGU (Tsekhmistrenko et al., 2019) were quite intriguing: while the plume conduit can continuously be followed down to the LLSVP in the deep mantle, the conduit is not as narrow and not nearly as vertical as classically expected!

Therefore I think it is quite safe to say that we have not yet heard the last of the Réunion hotspot and I hope that the next time you hear this name, maybe you will remember it as a rather remarkable spot on our planet…

 

Ashwal et al. (2017), Archaean zircons in Miocene oceanic hotspot rocks establish ancient continental crust beneath Mauritius, Nat. Commun., 8, 14,086, doi: 10.1038/ncomms14086.

Bredow, E. et al. (2017), How plume-ridge interaction shapes the crustal thickness pattern of the Réunion hotspot track, Geochem. Geophys. Geosyst., 18, doi:10.1002/2017GC006875.

Courtillot, V. E. and P. R. Renne (2003), On the ages of flood basalt events, C. R. Geosci., 335(1), 113–140, doi: 10.1016/S1631-0713(03)00006-3.

French, S. W. and B. Romanowicz (2015), Broad plumes rooted at the base of the Earth’s mantle beneath major hotspots, Nature, 525, 95–99, doi: 10.1038/nature14876.

Mazzullo, A. et al. (2017), Anisotropic tomography around Réunion Island from Rayleigh waves Journal of Geophysical Research: Solid Earth, 122, doi: 10.1002/2017JB014354.

Montelli, R. et al. (2006), A catalogue of deep mantle plumes: New results from finite-frequency tomography, Geochem. Geophys. Geosyst., 7, Q11007, doi: 10.1029/2006GC001248.

Morgan, W. J. (1971), Convection plumes in the lower mantle, Nature, 230, 42–43, doi: 10.1038/230042a0.

Morgan, W. J. (1972), Deep mantle convection plumes and plate motions, AAPG bulletin, 56(2), 203–213.

Morgan, W. J. (1978), Rodriguez, Darwin, Amsterdam, ..., A second type of Hotspot Island, J. Geophys. Res., 83(B11), 5355–5360, doi: 10.1029/JB083iB11p05355.

Richards, M. A. et al. (1989), Flood Basalts and Hot-Spot Tracks: Plume Heads and Tails, Science, 246, 103–107, doi: 10.1126/science.246.4926.103.

Scholz, J.-R. et al. (2018), SKS splitting in the Western Indian Ocean from land and seafloor seismometers: Plume, plate and ridge signatures, Earth Planet. Sci. Lett., Volume 498, 169-184, doi: 10.1016/j.epsl.2018.06.033.

Stähler, S. C. et al. (2016), Performance report of the RHUM-RUM ocean bottom seismometer network around La Réunion, western Indian Ocean, Adv. Geosci., 41, 43-63, doi: 10.5194/adgeo-41-43-2016.

Torsvik, T. H. et al. (2013), A Precambrian microcontinent in the Indian Ocean, Nat. Geosci., 6(3), 223–227, doi: 10.1038/ngeo1736.

Tsekhmistrenko, M. et al. (2019), Deep mantle upwelling under Réunion hotspot and the western Indian Ocean from P- and S-wave tomography, Geophysical Research Abstracts, Vol. 21, EGU2019-9447, EGU GA 2019.

Remarkable Regions – The Kenya Rift

Remarkable Regions – The Kenya Rift

Every 8 weeks we turn our attention to a Remarkable Region that deserves a spot in the scientific limelight. After looking at several convergent plate boundaries, this week the focus lies on part of a nascent divergent plate boundary: the Kenya Rift. The post is by postdoctoral researcher Anne Glerum of GFZ Potsdam.

Of course an active continental rift is worthy of the title “Remarkable Region”. And naturally I consider my own research area highly interesting. But after seeing it up-close and personal on a recent 10-day trip organized by the University of Potsdam, Roma Tre and the University of Nairobi (stay tuned for the travel log, or read that of the University of Potsdam), I must say, the Kenya Rift is a truly beautiful and fascinating region.

Figure 1. Topography (Amante and Eakins 2009) and kinematic plate boundaries (Sarah D. Stamps based on Bird 2003) of the East African Rift System (EARS). Plate boundary colors schematically indicate the western and eastern branches of the EARS.

Constituting one segment of the 5000 km long East African Rift System (EARS, Fig. 1), the Kenya Rift is host to an amazing landscape, wildlife and people, all of which somehow tie back to continental rifting processes. Although the youngest rifting phase in Kenya commenced in the Miocene, the east African region as a whole has been shaped by rifting episodes since Permian times (Bosworth and Morley 1994). The present active rift system runs from the Afar region in the north all the way south to Mozambique and is split into a western and an eastern branch that run around the Archean Tanzanian Craton (Chorowitz 2005, see Fig. 1). Generally speaking, the western branch is more seismically active, but deprived of magmatism, compared to the eastern branch, of which the Kenya Rift is part (Chorowitz 2005). Three processes characterize the EARS (Burke 1996) as well as the Kenya Rift specifically: normal faulting, volcanism and uplift.

Uplift

The Tanzanian Craton together with the enveloping western and eastern EARS branches constitutes the broad, uplifted area coined the East African Plateau (~1200 m elevation, Strecker 1991; Simiyu and Keller 1997, Fig. 2). The onset of uplift of this plateau can be constrained to the Early Miocene with the help of one of the longest phonolitic lava flows on Earth (> 300 km, Wichura et al. 2010; 2011) and a whale that stranded inland 17 Ma (and was only recently found again after going missing for 30 years, Wichura et al. 2015). Plume-lithosphere interaction is thought responsible for the uplift (e.g. Wichura et al. 2010), although there is disagreement about the continuity of the low seismic velocity anomalies seen in the east African upper mantle and whether they are connected to the lower mantle. For example Ebinger and Sleep (1998), Hansen et al. (2012), Sun et al. (2017) and Torres Acosta et al. (2015) advocate for one East African superplume, while Pik et al. (2006) distinguish separate lower and upper mantle plumes and Davis and Sack (2002) and Halldórsson et al. (2014) consider a lower mantle plume splitting in the upper mantle.

Figure 2. Topography (Amante and Eakins 2009) and fault traces (GEM) of the central EARS. Triangles indicate off-rift volcanoes, dotted grey lines the three segments of the Kenya Rift.

Magmatism and volcanism

The northward motion of Africa over this hot mantle anomaly has been thought the cause of a north-to-south younging trend in the age of the ensuing EARS volcanism and rifting (e.g. Ebinger and Sleep 1998; George et al. 1998; Nyblade and Brazier 2002), although more recent studies arrive at a more spatially disparate and diachronous rifting evolution (Torres Acosta et al. 2015 and references therein). In general, massive emplacement of flood-phonolites preceded the onset of rifting in Kenya around 15 Ma (Torres Acosta et al. 2015). With ongoing rifting, and localization of faulting towards the rift axis, volcanism also migrated towards the center of the rift. Since the Miocene, massive amounts of volcanics have thus been emplaced (144,000-230,000 km3, MacDonald 1994; Wichura et al. 2011). Moreover, dyking also accommodated a significant part of the extension, with 22 to 26 % of the crust in the rift valley being composed of dykes (MacDonald 2012). Not surprisingly, the highlands directly around the rift valley, the Kenya Dome (Fig. 2) formed through a combination of volcanism and uplift (Davis and Slack 2002) with elevations of up to 1900 m.

The composition of rift magmatism is bimodal, showing phonolites and trachytes on the one side and nephelinites and basalt on the other, predominantly resulting from fractional crystallization of a basaltic source. The low viscosity of these magmas allows the young volcanoes in the volcano-tectonic axis to reach significant heights (see Fig. 3; MacDonald 2012). The most impressive volcanoes are to be found outside of the rift however (Fig. 2), with Mnt. Elgon reaching 4321 m and Africa’s highest mountains Mnt. Kenya and Mnt. Kilimanjaro reaching up to 5200 m and 5964 m, respectively (Chorowitz 2005).

Figure 3. View on the crater rim of the 400 ky old Mnt. Longonot volcano in the tectono-magmatic rift axis, at 2560 m asl. Courtesy of Corinna Kallich, GFZ Potsdam.

Normal faulting

The Kenya rift itself is composed of 3 asymmetric segments, distinguished by sharp changes in their orientation (Chorowitz 2005, Fig. 2). The 2300-3000 m high Elgeyo, Mau and Nguruman escarpments result from the steep Miocene east-dipping border faults in the west, while the antithetic border faults on the eastern side formed later during the Pliocene (Strecker et al. 1990). The older border faults formed along preexisting foliation generated by the Mozambique Belt orogeny in the late Proterozoic (Shackleton 1993; Hetzel and Strecker 1994). A change in strike of this foliation from NNE in the northern and southern Kenya rifts to NW determined the change in orientation in the central Kenya rift (Strecker et al. 1990). Consequently, different generations of faults in the northern and southern rift segments run parallel, while in the central segment, the Pleistocene change in extension direction from ENE-WSW/E-W to the present-day WNW-ESE/NW-SE directed extension results in obliquely reactivated border faults and younger, en echelon arranged left-stepping NNE-striking fault zones along the rift axis (Strecker et al. 1990). Extension is transferred between the different zones by coeval normal and strike-slip faulting or dense sets of normal faults.

Figure 4. View of lake Magadi and the Nguruman escarpment. Lake Magadi is a saline, alkaline lake, commercially mined for trona. Courtesy of Corinna Kallich, GFZ Potsdam.

Human evolution

The uplift, volcanism and normal faulting together have set the stage for human and animal evolution. For example, the shift in hoofed mammals from eating predominantly woods to grazing species evidences that the large-scale uplift modified air circulation patterns resulting in aridification and savannah-expansion at the expense of forested areas (Sepulchre et al. 2006; Wichura et al. 2015). The rift basins enabled the formation of large lakes, which were subsequently compartmentalized by tectonic and volcanic morphological barriers (Fig. 4). On the short-term, lake coverage varied due to tectonically induced changes in catchment areas, drainage networks and outlets. Maslin et al. (2014) actually found a correlation between this ephemeral lake coverage and hominin diversity and dispersal. Lake highstands link with the emergence of new species and allowed the spread of hominins north and southward out of east Africa. Remarkable, or what!

References:
Amante, C. and Eakins B. W., 2009. NOAA Technical Memorandum NESDIS NGDC-24. National Geophysical Data Center, NOAA.
Bosworth, W. and Morley, C.K., 1994.  Tectonophysics 236, 93–115.
Burke, K., 1996. S. Afr. J. Geol. 99 (4), 339–409.
Chorowitz, J., 2005. J. Afr. Earth Sci. 43, 379-410.
Davis, P. M. and Slack, P. D. 2002. Geophys. Res. Lett. 29 (7), 1117.
Ebinger, C.J. and Sleep, N.H., 1998. Nature 395, 788-791.
George, R. et al., 1998.  Geology 26, 923–926.
Halldórsson, S. A. et al., 2014. Geophys. Res. Lett. 41, 2304–2311,
Hansen, S. E. et al., 2012.  Earth Planet. Sc. Lett. 319-320, 23-34.
Hetzel, R., Strecker, M.R., 1994. J. Struct. Geol. 16, 189–201.
Macdonald, R. et al., 1994a. J. Volcanol. Geoth. Res. 60, 301–325.
Macdonald, R., et al., 1994b. J. Geol. Soc. London 151, 879–888.
MacDonald, R., 2012. Lithos 152, 11-22.
Maslin, M. A. et al., 2014. Quaternary Sci. Rev. 101, 1-17.
Nyblade, A. A. and Brazier, R. A., 2002. Geology 30 (8), 755-758.
Pik, R. et al., 2006. Chem. Geol. 266, 100-114.
Sepulchre, P. et al., 2006. Science, 1419-1423.
Shackleton, R.M., 1993. Geological Society, London, Special Publications 76, 345–362.
Simiyu, S.M., Keller, G.R., 1997. Tectonophysics 278, 291–313.
Strecker, M., 1991. Das zentrale und südliche Kenia-rift unter besonderer berücksichtigung der neotektonischen entwicklung, habilitation, Universität Fridericiana.
Sun, M. et al., 2017.  Geophys. Res. Lett. 44, 12,116–12,124.
Torres Acosta, V. et al., 2015. Tectonics 34, 2367–2386.
Wichura, H. et al., 2010. Geology 38 (6), 543–546.
Wichura, H. et al , 2011. The Formation and Evolution of Africa: A Synopsis of 3.8 Ga of Earth History, eds. D. J. J. Van Hinsbergen, S. J. H. Buiter, T. H. Torsvik, C. and Gaina, S. J.
Wichura, H. et al., 2015. P. Natl. Acad. Sci. USA 112 (13), 3910-3915.

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].

References:
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.

 

References
Bemis, S. P., and W. K. Wallace (2007), Neotectonic framework of the north-central Alaska Range foothills, Geological Society of America Special Papers, 431, 549–572.
Bemis, S. P., R. J. Weldon, and G. A. Carver (2015), Slip partitioning along a continuously curved fault: Quaternary geologic controls on Denali fault system slip partitioning, growth of the Alaska Range, and the tectonics of south-central Alaska, Lithosphere, 7 (3), 235–246.
Bird, P. (1988), Formation of the Rocky Mountains, Western United States: A continuum computer model, Science, 239 (4847), 1501–1507.
Bird, P. (2003), An updated digital model of plate boundaries, Geochemistry, Geophysics, Geosystems, 4 (3).
Brocher, T. M., G. S. Fuis, M. A. Fisher, G. Plafker, Moses, M. J., J. J. Taber, and N. I. Christensen (1994), Mapping the megathrust beneath the northern gulf of alaska using wideangle seismic data, Journal of Geophysical Research: Solid Earth, 99 (B6), 11,663– 11,985.
Bruns, T. R. (1983), Model for the origin of the Yakutat block, an accreting terrane in the northern Gulf of Alaska, Geology, 11 (12), 718–721.
Castillo, P. R. (2006), An overview of adakite petrogenesis, Chinese Science Bulletin, 51 (3), 257–268.
Castillo, P. R. (2012), Adakite petrogenesis, Lithos, 134, 304–316.
Chapman, J. B., T. L. Pavlis, S. Gulick, A. Berger, L. Lowe, J. Spotila, R. Bruhn, M. Vorkink, P. Koons, A. Barker, et al. (2008), Neotectonics of the Yakutat collision: Changes in defor- mation driven by mass redistribution, Active Tectonics and Seismic Potential of Alaska, Geophys. Monogr. Ser, 179, 65–81.
Christeson, G. L., H. J. A. Gulick, P. S. ad Van Avendonk, L. L. Worthington, R. S. Reece, and T. L. Pavlis (2010), The Yakutat terrane: Dramatic change in crustal thickness across the Transition fault, Alaska, Geology, 38 (10), 895–898.
Chuang, L., M. Bostock, A. Wech, and A. Plourde (2017), Plateau subduction, intraslab seismicity, and the Denali (Alaska) volcanic gap, Geology, pp. G38,867–1.
Cloos, M. (1993), Lithospheric bouyancy and collisional orogenesis: Subduction of oceanic plateaus, continental margins, island arcs, spreading ridges, and seamounts, Geological Society of America Bulletin, 105 (715-737).
Cross, R. S., and J. T. Freymueller (2008), Plate coupling variation and block translation in the Andreanof segment of the Aleutian arc determined by subduction zone modeling using GPS data, Geophysical Research Letters, 34 (6).
Defant, M. J., and M. S. Drummond (1990), Derivation of some modern arc magmas by melting of young subducted lithosphere, Nature, 347 (6294), 662–665.
DeMets, C., and T. H. Dixon (1999), New kinematic models for Pacific-North America motion from 3 Ma to present, I: Evidence for steady motion and biases in the NUVEL-1A Model, Geophysical Research Letters, 26 (13), 1921–1924.
Eberhart-Phillips, D., D. H. Christensen, T. M. Brocher, R. Hansen, N. A. Ruppert, P. J. Haeussler, and G. A. Abers (2006), Imaging the transition from Aleutian subduction to Yakutat collision in central Alaska, with local earthquakes and active source data, Journal of Geophysical Research: Solid Earth, 111 (B11).
Elliott, J., J. T. Freymueller, and C. F. Larsen (2013), Active tectonics of the St. Elias orogen, Alaska, observed with GPS measurements, Journal of Geophysical Research: Solid Earth, 118 (10), 5625–5642.
Enkelmann, E., P. O. Koons, T. L. Pavlis, B. Hallet, A. Barker, J. Elliott, J. I. Garver, S. P. Gulick, R. M. Headley, G. L. Pavlis, et al. (2015), Cooperation among tectonic and surface processes in the St. Elias Range, Earth’s highest coastal mountains, Geophysical Research Letters, 42 (14), 5838–5846.
Ferris, A., G. A. Abers, D. H. Christensen, and E. Veenstra (2003), High resolution image of the subducted Pacific (?) plate beneath central Alaska, 50–150 km depth, Earth and Planetary Science Letters, 214 (3), 575–588.
Fitzgerald, P. G., R. B. Sorkhabi, T. F. Redfield, and E. Stump (1995), Uplift and denuda- tion of the central Alaska Range: A case study in the use of apatite fission track ther- mochronology to determine absolute uplift parameters, Journal of Geophysical Research: Solid Earth, 100 (B10), 20,175–20,191.
Freymueller, J. T., H. Woodard, S. C. Cohen, R. Cross, J. Elliott, C. F. Larsen, S. Hreins- dottir, and C. Zweck (2008), Active deformation processes in Alaska, based on 15 years of GPS measurements, Active tectonics and seismic potential of Alaska, 179, 1–42.
Fuis, G. S., T. E. Moore, G. Plafker, T. M. Brocher, M. A. Fisher, W. D. Mooney, W. Nodle- berg, R. Page, B. Beaudoin, N. I. Christensen, A. Levander, W. Lutter, R. Saltus, and
N. A. Ruppert (2008), Trans-Alaska Crustal Transect and continental evolution involving subduction underplating and synchronous foreland thrusting, Geology, 36 (3), 267–270.
Haeussler, P. J. (2008), An Overview of the Neotectonics of Interior Alaska: Far-Field Defor- mation From the Yakutat Microplate Collision, American Geophysical Union, pp. 86–108.
Haeussler, P. J., A. Matmon, D. P. Schwartz, and G. G. Seitz (2017), Neotectonics of interior alaska and the late quaternary slip rate along the denali fault system, Geosphere.
Haynie, K., and M. A. Jadamec (2017), Tectonic drivers of the Wrangell block: Insights on forearc sliver processes from 3D geodynamic models of Alaska, Tectonics, 36 (7), 1180– 1206.
Jadamec, M., and M. Billen (2010), Reconciling surface plate motions with rapid three- dimensional mantle flow around a slab edge, Nature, 465 (7296), 338–341.
Jadamec, M. A., and M. I. Billen (2012), The role of rheology and slab shape on rapid mantle flow: Three-dimensional numerical models of the Alaska slab edge, Journal of Geophysical Research: Solid Earth, 117 (B2).
Jadamec, M. A., M. I. Billen, and S. M. Roeske (2013), Three-dimensional numerical models of flat slab subduction and the Denali fault driving deformation in south-central Alaska, Earth and Planetary Science Letters, 376, 29–42.
Kerr, A. C. (2003), Oceanic plateaus, The Crust, Treaste on Geochemistry, 3, 537–565. Lallemand, S., A. Heuret, and D. Boutelier (2005), On the relationships between slab dip, back-arc stress, upper plate absolute motion, and crustal nature in subduction zones, Geochemistry Geophysics Geosystems, 6 (9).
Lease, R. O., P. J. Haeussler, and P. O'Sullivan (2016), Changing exhumation patterns during Cenozoic growth and glaciation of the Alaska Range: Insight from detrital geo-and thermo-chronology, Tectonics.
Marechal, A., S. Mazzotti, J. L. Elliot, J. T. Freymueller, and M. Schmidt (2015), Indentor- corner tectonics in the Yakutat-St. Elias collision constrained by GPS, Journal of Geo- physical Research: Solid Earth.
Martin-Short, R., R. M. Allen, and I. D. Bastow (2016), Subduction geometry beneath south-central Alaska and its relationship to volcanism, Geophysical Research Letters, doi: 10.1002/2016GL070580, 2016GL070580.
Mazzotti, S., and R. Hyndman (2002), Yakutat collision and strain transfer across the north- ern Canadian Cordillera, Geology, 30 (6), 495–498.
Morra, G., K. Regenauer-Lieb, and D. Giardini (2006), Curvature of oceanic arcs, Geology, 34 (10), 877–880.
Müller, R. D., M. Sdrolias, C. Gaina, and W. R. Roest (2008), Age, spreading rates, and spreading asymmetry of the world’s ocean crust, Geochemistry, Geophysics, Geosystems, 9 (4).
Page, R., C. Stephens, and J. Lahr (1989), Seismicity of the Wrangell and Aleutian Wadati- Benioff zones and the north American plate along the Trans-Alaska Crustal Transect, Chugach Mountains and Copper River Basin, Southern Alaska, Journal of Geophysical Research, 94 (B11), 16,059–16,082.
Peacock, S. M., T. Rushmer, and A. B. Thompson (1994), Partial melting of subducting oceanic crust, Earth and planetary science letters, 121 (1), 227–244.
Piromallo, C., T. Becker, F. Funiciello, and C. Faccenna (2006), Three-dimensional instan- taneous mantle flow induced by subduction, Geophysical Research Letters, 33 (8).
Plafker, G., J. C. Moore, and G. R. Winkler (1994a), Geology of the southern Alaska margin, The Geology of North America, The Geology of Alaska, G-1.
Plafker, G., L. M. Gilpin, and J. C. Lahr (1994b), Neotectonic map of Alaska, The Geology of North America, 1.
Preece, S. J., and W. K. Hart (2004), Geochemical variations in the <5 Ma Wrangell Volcanic Field, Alaska: implications for the magmatic and tectonic development of a complex continental arc system, Tectonophysics, 392 (1), 165–191.
Ratchkovski, N., and R. Hansen (2002a), New evidence for segmentation of the Alaska subduction zone, Bulletin Of The Seismological Society Of America, 92 (5), 1754–1765.
Ratchkovski, N. A., and R. A. Hansen (2002b), New Constraints on Tectonics of Interior Alaska: Earthquake Locations, Source Mechanisms, and Stress Regime, Bulletin Of The Seismological Society Of America, 92 (3), 998–1014.
Ribeiro, J. M., R. C. Maury, and M. Gr´egoire (2016), Are Adakites Slab Melts or High-pressure Fractionated Mantle Melts?, Journal of Petrology, 57 (5), 839, doi: 10.1093/petrology/egw023.
Richter, D., J. G. Smith, M. Lanphere, G. Dalrymple, B. Reed, and N. Shew (1990), Age and progression of volcanism, wrangell volcanic field, alaska, Bulletin of Volcanology, 53 (1), 29–44.
Rondenay, S., L. G. Mont´esi, and G. A. Abers (2010), New geophysical insight into the origin of the Denali volcanic gap, Geophysical Journal International, 182 (2), 613–630.
Schellart, W. (2004), Kinematics of subduction and subduction-induced flow in the upper mantle, Journal of Geophysical Research: Solid Earth, 109 (B7).
Smith, W. H., and D. T. Sandwell (1997), Global sea floor topography from satellite altimetry and ship depth soundings, Science, 277 (5334), 1956–1962.
Stegman, D., J. Freeman, W. Schellart, L. Moresi, and D. May (2006), Influence of trench width on subduction hinge retreat rates in 3-D models of slab rollback, Geochemistry, Geophysics, Geosystems, 7 (3).
Strak, V., and W. P. Schellart (2014), Evolution of 3-D subduction-induced mantle flow around lateral slab edges in analogue models of free subduction analysed by stereoscopic particle image velocimetry technique, Earth and Planetary Science Letters, 403, 368–379.
Waldien, T., S. M. Roeske, J. A. Benowitz, W. K. Allen, and K. D. Ridgway (2015), Neogene exhumation in the eastern Alaska Range and its relationship to splay fault activity in the Denali fault system, AGU Fall Meeting Abstracts.
Wang, Y., and C. Tape (2014), Seismic velocity structure and ansotropy of the Alaska subduction zone based on surface wave tomography, Journal of Geophysical Research: Solid Earth, 119, 8845–8865.
Worthington, L. L., S. P. Gulick, and T. L. Pavlis (2008), Identifying active structures in the Kayak Island and Pamplona zones: Implications for offshore tectonics of the Yakutat Microplate, Gulf of Alaska, Active tectonics and seismic potential of Alaska: American Geophysical Union Geophysical Monograph, 179, 257–268.
Worthington, L. L., H. Van Avendonk, S. Gulick, G. L. Christeson, and T. L. Pavlis (2012), Crustal structure of the Yakutat terrane and the evolution of subduction and collision in southern Alaska, Journal of Geophysical Research: Solid Earth, 117 (B1).