An Apple Pie From Scratch, Part Va: Tectonics: Construcing a Plate Tectonic History

USGS

It’s easy to imagine that, once a planet forms, its interior becomes a sort of independent realm, governed by its own laws and ignorant of the cosmos outside. In this model, the crust and its machinations are analogous to the thin scum that forms on a pot of soup; inconsequential byproducts, with no bearing on the whole. 

But in reality, a better metaphor for the evolution of a planet’s interior might be a stage play: There’s a lot going on behind the scenes, and for a complex production the stagehands may outnumber the actors. But what happens on stage drives everything else, and a perceptive audience member can infer the nature of the whole beast from what they are shown.

Crustal tectonics is not a sideshow. It is an engine converting heat from the interior into motion, and what happens on a planet’s surface can affect the behavior of the interior as deep as the core. As a planet ages and cools, this engine operates in different modes which can have a profound effect both on the landforms and conditions on the surface and on the hospitability of the planet to complex life.

• Earth Today 
• Patterns of Plate Motion 
• Convergent-Continent-Divergent 
• Convergent-Continent-Convergent 
• Convergent-Divergent-Convergent 
• Convergent-Divergent-Continent
•  Simulating Plate Tectonics
• Let Someone Else Do It
• Do It Myself: Quick and Dirty
• Do It Myself: Obsessively Detailed
• In Summary
• Notes

 Back to Part IVc

Earth Today

To start off, a quick tour: On average, the Earth’s radius is 6370 kilometers, but only the top 35 km is the crust (varying locally between 5 km at young sections of ocean and 70 km at old sections of continent). Underneath, the mantle stretches for over 2800 km to the boundary with the core.

Kelvinsong, Wikimedia

However, the distinction between the crust and mantle is based on chemical composition, as evidenced by the behavior of seismic waves passing through different materials. In terms of physical properties and motion, the uppermost region of the mantle is rigidly attached to crust and moves with the tectonic plates and so is often grouped with it in the lithosphere, 80 km thick (again, varying locally between 5 and 200 km). Below that, the asthenosphere stretching to 220 km depth is more ductile and fluid (though not liquid) and not attached to the plates, but does tend to be dragged along with them. The lower mantle—the mesosphere—stretches the rest of the way to the core-mantle boundary at 2980 km depth, is more uniform, and is largely composed of bridgmanite (MgSiO₃). Though it is at high temperature and can act as a fluid at large timescales (by continuously deforming), the high pressure keeps it mostly solid throughout.

Below that, the core is composed largely of a 9:1 mix of iron and nickel, with around 3% lighter elements (mostly oxygen and sulfur). The outer core, stretching to 5150 km depth, is liquid in spite of the high pressure thanks to a temperature of 4000 to 5400 K (hotter with depth). Though the inner core is even hotter—as high as 6000 K—it is under even higher pressure, around 3,500,000 atm, such that it is forced back into a solid.

As mentioned in the last post, most new heat is generated in the mantle and crust by decaying radioactive isotopes, but the core retains a good deal of primordial heat and is the best-insulated part of the planet, so has the highest temperature. As it cools, iron from the liquid outer core crystallizes onto the surface of the inner core. The liquid just outside the inner core thus becomes iron-depleted and so lighter and more buoyant than the overlaying liquid. It rises to the core-mantle boundary and there loses its heat to the cooler mantle, gains back its iron by mixing with surrounding liquid, and so becomes less buoyant and sinks back down towards the inner core, thus forming a convection cell. The liquid metal is conductive, and so this motion—along with the Earth’s rotation—forms the planet’s magnetic field.

"Advection" is transfer of heat by magma reaching the surface. Bkilli 1, Wikimedia.

The mantle, now heated by the core, convects as well, though much slower given its high viscosity. Its own radiogenic heat strengthens this current, and so once the plume of rock reaches the top of the mantle it breaks through the crust and reaches the surface. As the pressure drops the rock melts, and lava flows out of the seam opened in the crust. It quickly cools and solidifies, and is pushed aside by more lava bursting out and piling up around the seam. A ridge of new rock forms straddling this seam; the mid-ocean ridges we can see in many of our oceans today.

As this fresh rock is pushed further from the ridge, it cools and becomes denser, causing it to “float” lower on the underlying mantle, like a ship laden with cargo. This sinking helps to pull more material away from the ridge, and it’ll all move together as one tectonic plate. Eventually the rock will encounter another tectonic plate, and if that other plate is younger and less dense, then the rock will be pushed below it, forming a subduction zone. The rock sinks into the mantle, adding more pull to keep the rest of the plate behind it moving. It sinks lower, breaking up and mixing into the mantle as it goes, and eventually reaches the core-mantle boundary, closing the loop of mantle convection.

Surachit, Wikimedia/USGS

I’ve been emphasizing the pushing force from the mantle plume here to keep a narrative thread, but in reality the relative importance of “ridge-push” from mid-ocean ridges and “slab-pull” from subduction zones is an ongoing debate, and the current consensus leans towards the latter being more important. I’ll explain more in a moment, but the key point is that subduction causes ridge spreading, not the latter, and the mantle convection cell exists only so long as surface conditions allow it and the plate keeps moving.

What might stop a tectonic plate moving? Well, if the original seam of lava bursting through the crust is underwater—as is usually the case—then the rock forming at the mid-ocean ridge will have some water mixed in with it. As it travels from mid-ocean ridge to subduction zone, it will also tend to accrue layers of sediment from material settling out of the water. These sediments will trap more water, and some—like carbonate minerals—will themselves be formed of other volatiles. As the plate subducts and warms up in the mantle, these volatiles will melt out of the rock and rise upwards, puncturing the overlaying plate and bursting through the surface as volcanoes. The volatiles themselves will mostly be released as gasses, but as they melt and rise, they bring some of the surrounding silicate rock with them. Heavy elements like iron will tend to sink out of the mix as it rises, and so once this magma reaches the surface, it will form iron-depleted andesitic rock that is less dense than the basaltic rock formed at mid-ocean ridges. Islands of this material will pile up on the overlaying plate along the subduction zone, forming an island arc.

Given enough time, the individual volcanoes can produce enough rock to form a plate of their own; a section of continental crust, as opposed to the oceanic crust that forms at mid-ocean ridges. This lighter and thicker crust floats higher on the mantle, causing much of its surface to rise above the oceans.

Note that continental crust extends further down into the mantle as well. USGS

Once formed, continents grow by various mountain-forming events called orogenies. When a continent encounters plates of oceanic crust, the oceanic crust always subducts under the continental crust. As it does, more subduction volcanism causes more rock to build up on the continent along the plate boundary. If the oceanic plate already carries some volcanic islands formed of andesitic rock on it, these will collide with the continent and fuse with it, forcing some of the rock to fold up and form high mountain ranges.

Eventually subduction can pull two large continents together. They can continue to be pushed together for some time, forming a large plateau of folded rock, and one plate may even partially subduct under the other. But continental crust is too light to sink into the mantle, and so before long the plates will stop moving and fuse together. The thickened section of crust will be eroded from above and below, somewhat reducing the total amount of andesitic rock (some rock is also pulled from the overlaying plate at subduction zones), but this is a far slower loss process than subduction and cannot destroy a continent completely, so continents are largely permanent once formed. Sections of the continents are billions of years old, while the oldest oceanic crust—in the eastern Mediterranean—is no more than 280 million years old, and most of the oceans are far younger.

NOAA

Thus, our modern planet is a patchwork of old continents and young oceans. Most of the tectonic plates contain sections of both types of crust, but the sections of oceanic crust are ephemeral attachments to the more persistent continental cores.

The boundaries between plates can all be classified into 3 types:
 

domdomegg, Wikimedia

Divergent boundaries, where plates are spreading apart; mostly the mid-ocean ridges, such as in the Atlantic. But mantle plumes can also burst through continental crust, forcing the continent to spread apart and forming a rift valley, with volcanic ridgelines to either side. Before long, the continent is split apart, and a new ocean basin forms with a mid-ocean ridge. East Africa is undergoing rifting of this sort, though it’s unclear if this will ultimately split the continent apart or stop before that happens.

Red dots indicate earthquake epicenters. Note that in the continent-continent case (right) the continental crust will not completely subduct

Convergent boundaries, where plates are moving towards each other. These can be ocean-ocean subduction zones, such as the Marianas trench in the Eastern Pacific forming the Marianas island arc; ocean-continent subduction zones, such as the Peru-Chile trench off the western shore of South America, forming the Andes mountain range; or continent-continent collisions, such as the collision of India into Eurasia, forming the Himalayan plateau.

Transverse boundaries, where plates are moving in different directions past each other—or in the same direction but at different rates. Friction between these plates can cause earthquakes and land deformation, such as near the San Andreas fault in western North America, but no crust is being formed or subducted, and there is no significant volcanism.

In an idealized case such as the one I first described, a plate has a divergent boundary on one side and convergent on the other, such as is the case for the Nazca plate. But this need not always be a case. The African plate is surrounding mostly by divergent boundaries, meaning it is growing on all sides, while the Filipino plate is surrounded mostly by convergent subduction zones and will be completely consumed before long. The total area of the plates must, of course, remain constant, but the different motions of all the plates can allow area to effectively be traded between plates that are not in contact. Even where 2 plates are in contact, their complex shape or rotation of the plates can cause different types of boundaries to exist along the same border.

Major plates and the relative motion at their boundaries. USGS

No one plate can get too large, though; the older the oceans, the greater the chance they’ll subduct, and the larger the continent, the greater the chance that it will rift apart.

The key point here is that tectonic plates are not merely passengers on mantle convection. In a broad abstract sense the thermodynamics of the planet’s interior does provide the energy for plate motion, but once the plates start moving it is they that dictate the motion of rock through the mantle, not the other way around.

Patterns of Plate Motion

 The exact origins of plate tectonics are a bit murky, and a subject I’ll dig into more in the next post, but most researchers agree it has been ongoing in some form for at least 2 billion years. However it starts and whatever drives it, most of what we know about how plate tectonics works comes from observing Earth’s history, as interpreted through various surviving geological clues—common rock layers and fossil species across now-distant regions, records of shifts in the magnetic field and the orientation of continents relative to it, the motion of plates across stationary hotspots, and so on. Altogether, this gives us a near-complete record of how plates have moved and changed over the last 250 million years, since the supercontinent Pangea began to rift apart, and a fairly solid picture of motions going back 600 million years to the breakup of the supercontinent Pannotia. Past that there are multiple plausible models for the appearance and breakup of the supercontinent Rodinia about 1 billion years ago, decent indications for the assembly of supercontinent Nuna in some form 1.8 billion years ago, and a few vague hints at another assemblage around 2.5 billion years ago—perhaps another supercontinent we’d call Kenorland, perhaps multiple large continents in sequence.

Looking through all this data, there are a couple patterns that emerge. The most obvious is the supercontinent cycle. Over long periods the world’s continents have joined into a single landmass, rifted apart, and then joined together again several times, completing a cycle roughly every 750 million years. Of course we only have detailed records of a few of these cycles, and none are perfect: Rodinia started rifting before it was fully assembled, Pannotia appears oddly out of sequence and never included all major landmasses so you might just call it a stage in Rodinia’s breakup, and Pangea was spalling off microcontinents all throughout its tenure. But there is a definite trend, and it provides a convenient framework for constructing our own continents: Every continental world must be in some stage of the breakup, assembly, or tenure of a supercontinent. 

The exact mechanisms of the cycle are debated—particularly the role of the mantle—but we can at least say that the mantle helps in the rifting process. During its tenure, a supercontinent forms a large insulating layer above the mantle, and the trapped heat causes plumes of magma to form that eventually burst through the surface in the form of hotspots—isolated volcanoes—and flood volcanism—vast regions of volcanic activity and lava flows (it’s also possible the plumes have less to do with insulation and more with how patterns of subduction affect mantle convection, but at any rate these plumes consistently form under supercontinents). These plumes do not themselves cause rifting, but they form weak points in the crust which allows for subductions zones outside the continent to pull it apart. They also push the continental crust up slightly, creating a slight slope between the supercontinent core and surrounding subduction zones that the continents can slide down. 

Source

The rifting begins at triple junction, where rifting spreads out from weak points in 3 different directions. Eventually the spreading rifts from nearby triple junctions join together into a single zig-zag shaped rift, which you can still see in the overall shape of the modern Atlantic coasts. The third rifts from the triple junctions that don’t join with others become failed rifts and cease spreading, though they remain as weak points in the crust and can be reactivated later.

Rifts formed in the Pangea breakup. Note how failed rifts often become the cores of river valleys later. Source.

The rift zones will appear first as volcanically active rift valleys, but eventually split open into new ocean basins with mid-ocean ridges down their middles. Rifting often isn’t clean or simultaneous; Africa split from North America long before South America, and North America and Eurasia are still partially connected today over the Bering Strait.

USGS

Once the continents have rifted apart, they can reassemble into a new supercontinent in two primary ways. The simpler method is extroversion, where the new interior ocean basin continues to spread, the exterior ocean basin is fully subducted, and the continents collide more-or-less on the opposite side of the world from where they originally rifted apart, such that the exterior coasts of the old supercontinent become the interior of the new supercontinent. The alternative is introversion, where the interior ocean grows for a while, but new subduction zones form that cause the continents to stop and reverse course, closing the new basin and colliding along the same coasts that had previously rifted apart (though not necessarily with the exact same geometry). We’re not entirely sure what causes one or the other to happen, and often a particular supercontinent cycle can contain versions of both in separate ocean basins.

Earth in cross-section through different types of supercontinent cycles. Bradley 2011

Part of the mystery is uncertainty about how new subduction zones form. They’re unlikely to form directly between the interior ocean basin and attached continental coasts; though oceanic crust sinks as it ages, it also becomes stiffer and harder to break. More likely is that subduction begins along transverse boundaries, where the crust is already broken and two sections of oceanic crust of very different ages can come into direct contact. The geometry of linear spreading on a spherical surface requires that these boundaries must form somewhere within the new ocean, and once subduction starts there it can spread outwards along the continental coastlines and consume the entire interior ocean. Why these new subduction zones should come to dominate over the older subduction zones in the exterior ocean is an unresolved issue.

 As the supercontinent cycle is ongoing, there will be many local cases of rifting, collision, subduction, ridge spreading, basins opening and closing, and so on. I want to go through a few common geometries to clarify how they play out, but the essential rule throughout is that subduction is active and rifting and ridge spreading is passive. Subduction zones dictate the movement of plates, though the relative importance of different subduction zones seems to depend on mantle processes that aren’t fully understood. The best interpretation right now is that mantle convection forces the supercontinent cycle to occur, but subduction determines all the details.

Appearance of a convergent-continent-divergent system as it would appear on a map (above) and in profile through the crust (below). Arrows show plate motion and red/blue/green lines on the map show divergent/convergent/transverse boundaries. Image by me, using Paint.net.

Convergent-Continent-Divergent

This is the typical state of continents during an introversion cycle, with South America serving as a good example. The coast facing the direction of motion forms an active margin, with a subduction zone just offshore and a volcanic orogeny near the shore, forming a steep coastline. The leeward coast forms a passive margin, with well-eroded topography, vast coastal plains, and wide continental shelves.

The orogeny by the active margin is usually relatively thin, with a sharp transition to the interior lowlands, but in some cases a continent may experience flat-slab subduction, where a subducting plate doesn’t immediately dive into the mantle but slides underneath it for some distance before finally turning downwards.

EOS.org, modified from Finzel et al. 2016.

This causes uplift and deformation far into the interior of the continent, and either an end or motion inland of volcanism. This is how the Rockies in North America formed, though it's largely ceased now. It’s not totally clear why this happens, but it may be due to high speed of the continent relative to the subducted plate or subduction of a mid-ocean ridge.

Convergent-Continent-Convergent

This is the typical state of continents in the later stages of an extroversion cycle. There are no good current examples of this, but the eastern coast of Asia shows some of the features. Though both sides are active margins, there is still a clear direction of motion for the plate and different features on either side. The leading edge is much like the active margin in the previous example: an offshore subduction zone and coastal orogeny.

The trailing edge has a subduction zone and orogeny, but because of the continent’s motion the neighboring ocean is spreading faster than the subduction zone is consuming it. Thus the age of the subducting crust is increasing, meaning that it is becoming less buoyant. This leads to slab rollback; the ocean crust sinks further from the subduction zone, which has the effect of pulling the subduction zone away from the continent.

Niu 2014

The orogeny is pulled out with it, though because subduction is slower and the land is being stretched out it won’t form as high a mountain range as the leading edge, and can often be broken apart into an island arc. Between the volcanic arc and the continent interior the crust is thinned out, forming a back-arc basin. Fractures in the crust turn into hotspots, forming more volcanic islands. New oceanic crust can even form in the basin, but there is no mid-ocean ridge and the basin can never open up into a new ocean. The Sea of Japan is a typical back-arc basin. (Back-arc basins can form along the leading edge as well, but for different reasons and without significant thinning of the crust.)

Once the continent’s motion stops (e.g. if it collides with another continent) the trailing edge will be compressed and transition into a more typical active margin.

To be clear, the neighboring plates could be continents as well.

Convergent-Divergent-Convergent
This is the typical state of the exterior ocean during supercontinent tenure and breakup, with the Pacific and Nazca plates together being a good example. Such an ocean can remain stable for a long time so long as the subduction rate is symmetrical on both sides. But this is unlikely to be true forever, and so eventually the ridge will be subducted, leading to closure of the ocean.

In some cases there may not be a single ridge down the middle of an ocean, but 3 ridges meeting at a triple junction. Sometimes—not often, but sometimes—a new plate is formed in the middle of the triple junction, and then a triangle of ridges spreads outwards, growing the new plate at the expense of the old ones. Because this tends to push the ridges towards the edges of the plate, this can hasten the closure of the ocean. The Pacific plate formed this way, though it’s not yet clear if this will lead to its closure.

As the oceans close, island arcs formed over subductions zones will collide with the advancing continents, forming collisional orogenies along the coast—similar to volcanic orogenies in many ways, but tending to form higher and wider mountain ranges (though still not as wide as flat-slab subduction) with more internal deformation. The most recent example of this is along the coasts of Alaska and British Columbia, though this orogeny is largely concluded by now.

When the oceans finally close, the subduction zones will not close all at once, but may remain in fragments along complex coastlines. The continents may slow in some areas or even stop while trapped sections of ocean remain. The surviving sections can undergo slab rollback, closing some sections of ocean and widening others, leaving complex patterns of land and sea. The Mediterranean was formed in this fashion.

Where the oceans do completely close, the continents form one last collisional orogeny in between them. This has started in some sections of the Mediterranean (in the Alps and the Caucasus) and the Urals are a good example of the eventual result. These collision zones remain weak points in the crust, and can rift apart again later.

Convergent-Divergent-Continent 

This is the typical state of an interior ocean in the later stages of an extroversion cycle. The former Tethys ocean was an ideal example, with the current collision between India, Arabia, and Africa into Eurasia as the final result. Because ridge spreading is symmetrical and the crust is only being subducted on one side of the ocean, the ridge must move towards the subduction zone and eventually be subducted by it. This can lead to closure of the ocean, much as in the previous case, but in some cases the continents on either side of the subduction zone may be connected elsewhere, such that they can’t move relative to each other.

The subduction zone keeps pulling on the subducting plate, and so it has to break somewhere, and though it is thicker continental crust tends to break easier than oceanic crust. If the continent across the ocean has no internal rifts, then a section of that continent on the coast facing the subduction zone will be torn off, forming a microcontinent (or multiple along the coast) with a new mid-ocean ridge behind it.

The microcontinent will cross the ocean and collide with the continent or island arc on the far side of the subduction zone, forming a collision orogeny, but at that point the subduction zone can “jump” across the microcontinent—that is, a new subduction zone will form on the trailing side of the microcontinent after or even before the collision. That puts us right back where we started, with a subduction zone, ridge, and continent passive margin.

Eventually the continent—or a large section of it—must close the ocean basin and collide with the continent overlaying the subduction zone. Because the continent is part of the subducting plate, it will collide with the overlaying plate at high speed, and because the margin of the overlaying plate is composed of many pieces of former microcontinents, it is fairly weak. Thus the collision will produce a wide, high, intensely folded collision orogeny, such as the Himalayan Plateau we have today. Only such a large continent-continent collision can permanently stop a subduction zone.

Plate speed depends on a couple different factors. Subducting plates move fastest due to slab pull, and subduction even helps pull overlaying plates through slab suction. Ridge push has a secondary effect, which is stronger for plates spreading away from a former subcontinent thanks to the slight slope created by mantle plumes—but to be clear, a continent can only be moved by ridge push if there is subduction elsewhere helping it along. Countering these forces is drag from the mantle, which has a stronger effect on continents because they have deeper lithospheric “roots”. As points of reference:

• India, on a mostly-oceanic subducting plate, topped out at 20 centimeters/year in the late Cretaceous, and continues to move north at around 6 cm/yr.
• The Nazca plate, a subducting oceanic plate, is moving east at 10 cm/yr, and the Pacific plate west at around 8 cm/yr.
• North and South America, large continents with large active margins, are both moving west at 3 cm/yr.
• When it initially starting rifting from Africa, North America only moved at 1 cm/yr for the first 25 million years, before abruptly accelerating to 3.5 cm/yr in a 6-million-year-window.
• Eurasia, a huge continent with a small active margin, is only moving about 1.5 cm/yr. Despite being a “convergent” boundary, the eastern coast is actually spreading at 1-2 cm/yr to accommodate the motion of the Americas (which is to say, oceanic crust is still subducting but the boundary is moving away from the mid-ocean ridge), which is causing slab rollback and back-arc extension.
• Africa and Australia, continents surrounded by divergent boundaries, both move at less than 1 cm/yr.

1 cm/yr is about 0.9° of rotation around the Earth every 10 million years (i.e. 0.9° latitude or longitude on the equator). In other words:

Situation	Subducting Ocean	Recent Subduction Collision	Active Margin Continent	Passive Margin Continent
Plate speed (°/10 myr)	7.2-18	5.4	2.7	<0.9

For planets of other sizes, divide these values by the radius in Earth radii (presuming equivalent plate speed, which probably isn’t true but there’s no way to get a better estimate right now).

There appears to have been a general increase in average plate speed since the onset of plate tectonics, roughly doubling over the last 2 billion years, which may be associated with the period of the supercontinent cycle decreasing by the same factor, but at any one time overall variation will be larger than this general trend. Plates can accelerate very quickly when a new subduction zone is formed or a mid-ocean ridge is subducted.

Note that the speed of a plate’s movement is not necessarily the speed at which it’s boundaries are moving, save for in the case of a continent’s leading-edge active margin; an ocean plate can be moving at high speed but bounded by static convergent and divergent margins (though no boundary is totally still today).

The area of the continents is also increasing, but only slightly. The Earth gains an average of 0.6-0.9 cubic kilometers of continental crust per year—at an average thickness of 35 km, that’s 0.017-0.025 km²/year, or 17-26 thousand km²/million years, compared to a current continental crust area (including submerged continental shelves) of ~200 million km². That’s around 9-15% growth over the last billion years; not a lot, but worth bearing in mind (all of this assuming average continental crust thickness has remained constant, which might not be the case). This may go faster on larger planets because of square-cube scaling between the area of the crust and the volume of the mantle, but it's not clear.

Finally, sea levels tend to rise and fall relative to the continents with the supercontinent cycle. During a supercontinent’s tenure, the oceans are mostly composed of old, low-lying oceanic crust and the continents are lifted up by mantle plumes, combining to cause low sea levels. Once the continent breaks up, the old oceanic crust is replaced with younger, more buoyant crust and the continents move off the mantle plumes, so sea levels rise over the next hundred million years or so, until the oceanic crust ages and continents begin reassembling. Global climate and glaciations also affect sea level, but the relationship of those to tectonic events is harder to parse; we’ll tackle that in a later post.

Simulating Plate Tectonics

To get a map with a realistic tectonic history, there’s a few different methods; 3 that I'll share here today. I don’t suggest trying to retcon a tectonic history onto a map you’ve already made—sorry—or trying to just slap down plate boundaries at random. These boundaries are formed by the patterns of continental drift, not the reverse.

Let Someone Else Do It
Find a program to model the tectonic history for you. The best available one I know of right now—by a huge margin—is that in the alpha version of Songs of the Eons, a strategy game in development. The map generator simulates a tectonic and even glacial history for the planet, and the results are fairly impressive from a standpoint of geological detail and accuracy (though sadly it can’t quite seem to do fold-and-thrust belts properly, probably due to resolution limitations).

Elevation map produced by Songs of the Eons--from their website.

Some other map generators (e.g. Azgaar's) are fine for small regions, but I advise strongly against map generators that just rely on fractal cloud generation, without applying a lot of tectonic modification to it. They may produce features that appear “organic” at a glance, but they miss a lot of the nuance of real tectonic systems. Rather than a diverse world, you end up with pretty much the same ragged coasts, rolling hills, and gently rising highlands everywhere on the planet. Watch out also for generators that tend to make landmasses with high interiors gently sloping to low coasts. I’ve already mentioned that many orogenies are focused on coastlines, and often the distinction between highland and lowland is quite distinct.

Do It Myself: Quick and Dirty
If you want more control over how the resulting continents look, then we can try to make our own tectonic history by hand. But if you don’t want to spend too long on it, then we can use supercontinent cycle as a shortcut by working forward from the last supercontinent to work out a reasonable modern distribution of continents and plate boundaries. Most people seem to want a world roughly resembling modern Earth—late in the breakup of the last supercontinent—so I’ll start with one of those.

Basically, we just take a vaguely positioned supercontinent, slash it down the middle with a zig-zagging rift, and move the resulting fragments directly away.

My supercontinent. Red is the rift, arrows indicate motion of the continents.

Break apart the fragments a bit, rotate them, maybe crash a couple of them into each other, and we should be good.

The arrows within the continents show the rotation of each fragment.

We should end up with an interior ocean surrounded by passive margin coasts—that are vaguely mirrored on opposite sides—and an exterior ocean surrounded by active margins.

Blue lines are active margine, grey lines are passive margins.

If we follow modern Earth’s example we might designate the active margins on one side of the ocean as leading-edge, with smooth coastlines and high coastal ranges, and those on the other side as trailing-edge, with jagged coastlines and offshore island arcs. But we could also make them all leading edge, and in either case place a couple more subduction zones between the continents, and perhaps even breaking in between them into the interior ocean as has happened in the Caribbean on Earth.

Blue lines are leading-edge active margins, orange lines are trailing-edge active margins, purple lines are new subduction zones.

Finally in the continent interiors we can place high mountain ranges on active margin coasts and continent collision boundaries and low mountains randomly and along the interior coasts as remnants from older collisions.

Dark brown is new high mountain ranges, beige is old lower ranges

If we wanted a world during supercontinent assembly, we have to decide whether to do an extroversion or introversion cycle. For extroversion, we keep moving the continents out—rotating and crashing them into each other as they go—until they start meeting on the opposite side of the world—throwing in some trailing-edge active margins on the coasts of the interior ocean as the crust there ages and more subduction zones break into it from the sides.

Orange lines are new active margins, and brown areas new collisional mountain ranges. I haven't adjusted to coastlines to reflect this, just moved the landmasses.

For introversion, we place subduction zones in the interior ocean, but turn them into leading-edge active margins and leave the exterior coasts as trailing-edge active margins (this needn’t happen on all sides of the ocean, just one will do). Carry the continents back towards each other, but with some rotation as they’re pulled about by the new subduction zones; the coasts shouldn’t meet back up neatly in the same places they split

Blue lines are new leading-edge active margins, orange lines are former leading-edge active margins that have become trailing-edge. Again, coastlines aren't adjusted.

If we want a supercontinent, we just carry forward either case until the continents have all met back up and closed the oceans between them, leaving high mountain ranges along the sutures and active margins on most—or all—of the coasts. The supercontinent should still be moving somewhat relative to the oceans, so we’ll have a leading edge and trailing edge, and we can also have microcontinents moving around the edges if we don’t want all the land area in one piece.

New supercontinent, carried forward from the introversion model. Blue lines are leading-edge active margins and orange lines are trailing-edge, after the supercontinent has formed.

 

Do It Myself: Obsessively Detailed
If we really want a complete, detailed, and realistic history for the whole planet, then we need to fully simulate the history back not just from the last supercontinent, but all the way back to the supercontinent before it—around 500 million to 1 billion years of history. This is far enough back to cover the formation of almost all major geological features on Earth today.

And while we're doing that, we'll be tracking not just the movement of the continents, but all the plates, and the boundaries between them.

So we want to start with a supercontinent, break it apart, let it reassemble into a new supercontinent, and then break that apart and continue to the stage in the cycle we want to reach, all the time following the rules we established in the last section. As we do this, we’ll be recording the positions of a few key geological features:

Orogenies, distinguishing between Andean-type orogenies, where a subduction zone forms along the coast of a continent; Laramide-type orogenies, where a continent quickly subducts a mid-ocean ridge; Ural-type orogenies, where a continent collides with an island arc, microcontinent, or the active margin of another continent; and Himalayan-type orogenies, where a continent on a subducting plate collides with a continent on an overlaying plate. We shouldn’t have to worry much about Japan-type trailing edge subduction zones, as the affected regions should experience larger orogenies in short order—we just have to note the ones still active at the end.

For reference, here are the widths of modern examples of these orogeny types:

Location	Orogeny Type	Width (km)
Ecuador	Andean	80
Chile/Argentina	Andean	120
Peru	Andean	200
Chile/Bolivia	Laramide	750
US southwest	Laramide	1300
Russian Urals	Ural	50
Switzerland/Italy	Ural	180
British Columbia	Ural/Andean	600
Nepal/China	Himalayan	1200

Failed Rifts, (formally called aulacogens) as these may rift again later, and even if they don’t they tend to form low-laying regions that become the core of major river valleys later on, like the modern Mississippi.

Hotspots, which form local volcanic features and remain mostly (but not always) stationary relative to the mantle—Hawaii is a prime example. There is some evidence that these are mostly associated with sections of subducted crust sitting above the core-mantle boundary, but they can appear elsewhere. As a guideline, we’ll expect them to mostly appear either near the rifting center of a supercontinent, roughly on the opposite side of the planet, or in areas of weakened crust like trailing edge back-arc basins, but they’re also possible anywhere else. They’re small features that tend to be erased by other processes in the long run, so we don’t really have to worry about them until the last 100 million years or so of the sequence. Definitions vary but there are around 60 hotspots active today, though many are very small or appear at continent boundaries where we already expect there to be volcanic activity—so really 10-20 should be fine.

Large Igneous Provinces, which are massive regions of flood basalts created by giant mantle plumes. Iceland is the closest equivalent we have today, but it’s nowhere near the same scale (I’ll describe them and their effects in more detail when I discuss mass extinctions). They appear in roughly the same areas as hotspots, though with no particular association with trailing edges—LIPs are more consistently associated with large mantle plumes. The largest known—the Central Atlantic Magmatic Province—formed along the North America-Africa rift at the start of Pangea’s breakup. We’ll only concern ourselves with the largest examples, and also only those on the continents—though they appear frequently in the oceans as well, they don’t typically create islands or any other particularly notable feature. Again, definitions vary but there have been about a dozen large continental LIPs in the last half-billion years, maybe 5 of which were associated with the breakup of Pangea. Also note that LIPs often leave hotspots behind after they conclude.

To do all this, I’ll be using a program called GPlates, a free program designed to allow paleogeographers to map the motions of Earth’s plates in the past and does a pretty decent job with fictional plates as well. In principle everything I’m going to do here could also be achieved with a beach ball and some cut-out pieces of paper, but using software just seems easier for me. I would not recommend attempting this on a flat map.

I've now added a supplemental tutorial on using GPlates in a separate post; I still recommend the tutorial written up here as a secondary source. Refer to those for the specifics of using the program; what I want to lay out here is a more conceptual outline of how to simulate tectonics in semi-realistic way; at least, well enough to produce realistic "modern" geography.

Remember that you can click each of these images for a bigger version. All are created by me in GPlates and exported to Mollweide projections.

I’ll start out with an ideal supercontinent containing 10 major cratons (marked in grey). These are particular old sections of continental crust at the core of all of Earth’s continents that are thicker and tougher than the rest and so are rarely broken or even much deformed—so we’ll keep them intact through the exercise, and give them rather flat interiors at the end. Each craton has some accreted continental crust around it, and for now there are no additional bodies outside the supercontinent.

Overall this supercontinent occupies about 1/4 of the planet’s surface area in order to about match Earth’s land area around a billion years ago, but there’s no particular reason that should be the case for all worlds. We don’t need to worry about the plates in the oceans or any features of the continents aside from the placement of the cratons, because all of these will be near-completely erased and overwritten with new plates and features over the course of this exercise.

From here I'll move forward in 50-million-year steps, moving the continents, creating new oceans, and splitting and fusing plates as required by the rules laid out so far—though still with some space for artistic interpretation in the exact speed and direction of plate movement and placement of new rifts. Though we haven't marked out any boundaries in the oceans, we only need a few subduction zones to start out with and before long we'll have completely replaced the old oceans with new crust.

But first, a little bookkeeping: I’ll give each of the cratons a letter designation, and refer to any continents they form by the combination of the cratons in them; so this first landmass carries the admittedly awkward title of supercontinent ABCDEFGHIJ. As new ocean plates appear I’ll designate them with lowercase Roman numerals (i,ii, iii, iv, etc.) and microcontinents (sections of continental crust without a craton) I’ll designate with numbers (1, 2, 3, 4, etc.).

I also want to break up the sequence of events into geological periods, but with more memorable names. I’ll define the time starting with the breakup of this supercontinent as the Cuvieric Period.

I’m going to be taking a very “let’s see how it goes” approach, but just to keep everything organized I’m going to decide now that our first supercontinent cycle should be an extroversion cycle. Using my powers of prophecy, I can see that this means most of the action is going to occur on the other side of the planet, awkwardly on the edges of the above map—so for the sake of consistency I’m going to shift our view 105° west. I can also see that the entire process is going to take 850 million years.

Let’s get started. 

 850 mya: Early Cuvieric

We start off much as we did before, with a major rift bisecting the supercontinent (marked in red). But this time I’ve made sure to avoid the cratons, and mark out the third-arm failed rifts. I’ve also added subduction zones around the rim of the supercontinent to help pull it apart (marked in blue) and noted the Andean-type orogenies caused by them (marked in black).

Presumably there would be LIPs and hotspots associated with this rifting, but the landforms created won’t last 850 million years so we don’t have to worry about them this time around. It’s just something to bear in mind when we work out the history of life on this world. 

800 mya: Mid Cuvieric

The two supercontinent halves—ABCDEF and GHIJ—drift apart, leaving a mid-ocean ridge between them (marked in red, like the rifts) and transform boundaries on their edges (marked in green). New ocean crust forms between them (shaded in lighter blue than the Pre-Cuvieric ocean crust). They do not break apart evenly—on such a long border on a spherical planet that’s not really possible, but in this case especially the subduction zone on GHIJ’s north shore helps rotate the continents as they split.

Arrows indicate relative plate motion at boundaries, with length roughly correlating to speed. All these closeup maps are in orthographic projections; i.e., like looking at a globe.

But already, GHI is rifting away from J, with associated failed rifts—the main rift itself is an extension of a “failed” rift from the initial rifting event. Such multi-stage rifting appears to be typical of supercontinent breakups, with all the resulting continents moving away from a specific rifting center.

As the continents move, the subduction zones on their flanks churn away, and I’ve added volcanic island arcs along those in the oceans. I’ll continue to do so as new subduction zones appear, though to be clear these aren’t meant to be accurate depictions of the types of islands that would appear—just abstract representations that there are islands there.

750 mya: Late Cuvieric

The continents continue to drift apart, and now a triple-junction of mid-ocean ridges has formed between them. I’ll add one more rift to split ABCD from EF—again along a prior failed rift—and call it good so far as the supercontinent breakup is concerned.

But already there are other forces causing rifting: The subduction zone on GHI’s east coast has spread south, creating an enclosed ocean basin with a subduction zone on one side and no mid-ocean ridge to feed it. As such, a section of land—Microcontinent 1—is being torn away from shore opposite the subduction zone.

By this point, new ocean crust and motion of the continents away from the magma plume under the rifting center has caused a significant rise in sea level, and we can expect inland seas on many of the continents—though I won’t bother showing that here.

With the supercontinent rifting coming to an end and other tectonic processes starting to play out, we’ll call this the end of the Cuvieric and start another period, the Anningic.

700 mya: Early Anningic

The continents continue to rift apart, moving at increasing speed. EF and J—which is now crossing over the north pole—are speeding towards each other due to the subduction between them, and Microcontinent 1 is moving particularly fast towards a collision with the other shore of GHI. 

650 mya: Late Anningic

Microcontinent 1 collides with GHI at 660 mya, forming a continent-continent convergent boundary (marked in blue) and a Ural-type orogeny (marked in black). The subduction zone “jumps” across the thin landmass, consuming the new oceanic crust created behind 1 as it travelled.

EF is speeding towards its own collision with J, and the pull of J’s subduction zones is causing it to rotate clockwise. This compressed the oceanic crust near its north shore, and transform boundaries there become a new subduction zone that spreads down its east coast. The mid-ocean ridge to its east continues to spread, but now will drift west due to the subduction. This is a typical ocean introversion cycle, but at a local rather than global scale.

As EF drifts away from ABCD, the subduction zone on its west coast spreads as well. EF is still moving westward, even though it is subducting crust on it east side.

600 mya: Early Owenian

At 620 mya, EF finally collides with J. We’ll take this collision to conclude the Anningic and start another period, the Owenian.

As the continents collide and slow down, subduction pulls E and F apart, and the subduction zone on E’s east coast has consumed the mid-ocean ridge there and begins pulling a section of land, Microcontinent 3, away from J. To the south, E’s subduction zone spreads and links up with F’s subduction zone, which allows the whole assemblage—EFJ—to drift south.

Meanwhile, the subduction zone on GHI’s interior ocean has consumed the mid-ocean ridge there, and begins pulling another microcontinent—Microcontinent 2—away from the far shore. 

550 mya: Mid Owenian

The complicated EF-J collision continues. E and F have now both solidly collided with J, forming Ural-type orogenies, but a remaining section of subduction zone pulls a small landmass, Microcontinent 4, from E’s shore. In the east, 3 begins to collide with E.

In the south, Microcontinent 2 swiftly moves towards collision with GHI, but GHI itself is moving towards collision with ABCD.

At sea, the mid-ocean ridge between EFJ on one side and ABCD and GHI on the other has grown so broad that spreading along its whole length is no longer possible—a result of the geometry of spherical planets. A subduction zone forms along its eastern edge—north of GHI—which starts pulling GHI and EFJ together. EFJ has also consumed the mid-ocean ridge to its south, so we’ve solidly moved from supercontinent breakup to supercontinent assembly. Though we said it would be an extroversion cycle, as is typical it’s actually more of a mix: an extroversion between ABCD and GHI, and an introversion between EF and J and between EFJ and ABCDGHI

With continents over both poles, we can expect climate to be cooling—perhaps there’s even an ice age. This, along with aging of the new oceans, will pull sea level down from its late-Cuvieric high.

That said, there might be a brief warm spell associated with our first LIP in northeast GHI (marked in orange), as well as a possible extinction event.

500 mya: Late Owenian

The EF-J collision event has now largely concluded; E and F are solidly attached to J, and Microcontinents 3 and 4 have collided with E and J, respectively. The trapped sea between E and J would probably be filled in before long, but I’ll leave it in place as a weak point in the crust. I could also have the subduction zone “jump” these microcontinents and continue pulling chunks away—this is probably more realistic—but given that these regions will soon be involved in the collision with the southern continents, I’ll leave them for now.

As EFJ moves south and its island arcs collide with ABCD’s, the islands accrete to the overlaying plate and form a section of continental crust—eventually a whole new continent could be formed this way, but that won’t be the case here.

In the south, Microcontinent 2 collides with GHI at 530 mya, compressing the already-deformed lands of former Microcontinent 1. If this process continued uninterrupted, H might eventually be torn away and finally close the subduction zone—but again, that won’t be the case here.

Note here that we can infer the presence of an old mid-ocean ridge from before our simulation began due to relative plate motion around the pre-Cuvieric ocean.

450 mya: Early Huxleyic

ABCD finally collides with GHI at about 460 mya, closing the Owenian and starting the Huxleyic period. As with the EF-J collision, it’s a messy process, and subduction zones pull G—and its accreted microcontinents—away from HI.

Meanwhile, the oceanic crust formed at the start of the Cuvieric is now pushing 400 million years old. A subduction zone spreads down the west coast of GHI, eventually separating the continent from the oceanic crust and forming our first ocean-only plate, Ocean i, which continues to move east.

A subduction zone also spreads along EFJ’s eastern coast, so the continent is pulled in 2 directions. E splits from FJ, taking a large part of J’s former plate with it. FJ continues south towards the already-starting collision with the assembling supercontinent, and E drifts east out to sea.

Just for fun, I’ve also dropped a small LIP in FJ.

400 mya: Mid Huxleyic

The ABCD-HI collisions continues, with widespread mountain formation and land deformation, including a proper Himalayan-type orogeny between ABCD and G, where 2 microcontinents are sandwiched.

FJ also collides with the supercontinent’s north, forming an impressive orogeny of its own thanks to the trapped sections of former island arcs.

E continues east, though it’s rotating clockwise thanks to subduction between it and ABCD. It also experiences a massive LIP, potentially causing a large mass extinction.

At sea, the triple junction of mid ocean ridges spreads apart to create another plate of new oceanic crust, Ocean ii. Though it has no subduction zones of its own, this plate will spread at the expense of surrounding plates as they are subducted near the coasts. At this point, almost all remaining oceanic crust from before the Cuvieric has been subducted (save for a couple patches near the south pole and between the colliding continents), so we have a complete image of the tectonic forces at play in our world.

One consequence of this is that to keep Ocean i moving in the right direction to feed all the subduction zones on GHI’s west coast, it has to rotate clockwise and so start subducting part of the ocean connected to ABCD.

350 mya: Late Huxleyic

The ABCD-GHI collision concludes and FJ more solidly attaches to the assemblage. We can definitely say that we’ve entered the tenure of a new supercontinent, ABCDFGHIJ. We’ll mark the occasion by closing the Huxleyic and starting the Marshian period.

We could let the situation continue to play out until E joined the supercontinent, but just to make things a little more interesting we’ll have a minor rifting event that separates D and Microcontinent 5 from the supercontinent’s east coast.

As these new landmasses move east, the subduction zone on E’s east coast pulls the distant ocean crust west. The result is a new subduction zone splitting the 2 sections of crust apart, forming Ocean iii. Meanwhile, subduction zones spread along the north coasts of E and FJ, splitting them from Ocean iv. There is now a complete ring of subduction zones separating the assembling supercontinent from the surrounding superocean—which is exactly what should happen as the oceans age and the continents stop moving relative to each other.

As Ocean I moves south, the mid-ocean ridge between it and Ocean iv begins to subduct under the supercontinent’s west coast, so we’ll put a Laramide-type Orogeny there.

300 mya: Early Marshian

Internal mountain building has largely ceased on ABCDGHI, though it continues along subduction zones and a sliding collision in the south with Microcontinent 5. The whole landmass drifts slowly north and rotates clockwise. The vast interior will be fairly arid, with a large desert at the center.

D drifts north and E drifts south, accreting island arc crust along its south coast. A small section of the ocean crust attached to Microcontinent 5 gets pinched out between E and Ocean I, forming a small ocean plate, but it won’t last long enough to be worth giving a name.

Ocean ii continues to grow, and the very last of the pre-Cuvieric oceans have either been consumed by subduction or trapped within the supercontinent. Some small fragments will have been lifted onto the continents and survive there as ophiolites.

250 mya: Late Marshian

Microcontinent 5 impacts E, sandwiching some island arcs, and other island arcs from ocean i impact E’s south coast, creating a long, complex orogeny. The subduction zone “jumps” across plate 5, and E is pulled towards a collision with the supercontinent.

D continues north, and has a moderate LIP.

The mid-ocean ridge between Ocean iv and Oceans i, ii, and iii is now so broad that, as has happened before, it can no long spread across its entire length. Ocean iv begins subducting below Ocean i.

200 mya: Early Copian

E finally collides with ABCFGHIJ at about 220 mya, forming a vast Himalayan-type orogeny, and D begins colliding with the same landmass, so for short time we have a complete supercontinent ABCDEFGHIJ.

But not for long. A rift crosses the interior, and with it comes a massive LIP, flooding the region with lava. ACG in the south begins splitting from BFHIJ in the north. So ends the Marshian and begins the Copian period.

The northward motion of BFHIJ encourages the formation of a new subduction zone on its north coast to subduct the old crust inside the now-separated Ocean v. Ocean i is also rapidly subducting, and the subduction of its mid-ocean ridge with Ocean ii causes a Laramide-type orogeny along ACG’s west coast.

Rifting also breaks apart E. The new landmass contains much of the crust from former Microcontinent 3, back in the Owenian, so we’ll give it the same name. It may look like an odd snaking continent here, but in reality it would probably appear as a long archipelago like Indonesia.

Now that E has stopped, a subduction zone spreads along its north coast, separating it from Ocean vi. This subduction also pulls away a section of D, Microcontinent 6.

150 mya: Late Copian

The supercontinent has now solidly split in two, but like last time it doesn’t spread evenly; subduction zones in the east cause BFGHIJ to rotate clockwise and ACEG to rotate counterclockwise.

Also like last time, the rifting doesn’t conclude in one step. G begins rifting from ACE, with an associated LIP.

To BFHIJ’s north, Ocean i has subducted completely and Ocean v will soon follow. Their associated island arcs attach to BFHIJ’s coast, forming Ural-type orogenies. This is the typical fate of island arcs offshore of continent active margins.

In the east, D finishes colliding with BFHIJ and Microcontinent 6 speeds south towards ACE. In this process, another oceanic plate is cut off, Ocean vii. This whole eastern ocean is now a maze of subduction zones, island arcs, and mid-ocean ridges, but it’s all being inexorably pulled towards collisions with the continents.

In the south, Microcontinent 3 collides with G to its west and islands from Ocean vi to its east. Take special note here how odd this appears in the Mollweide projection, as opposed to the view over the south pole below; always keep a map's distortion in mind when working near the poles.

Finally, as we approach the end we should start adding hotspots (marked with purple dots)—some near rifting centers, some randomly across the globe. For now, though, just one left over from the Early Copian LIP.

100 mya: Early Andrewsian

At this point I’m trying to construct more of a continuous geological history rather than doing everything in 50-million year steps. As such, shortly after G rifts from ACE, CE rifts from A at 110 mya—closing the Copian and beginning the Andrewsian period.

But the supercontinent breakup isn’t over—an LIP within BDFHIJ foretells another rifting event.

Microcontinent 6 also collides with CE around this time, and the mid-ocean ridge behind it is quickly consumed. This pulls CE towards BDFHIJ, starting an introversion cycle.

Oceans vi and vii are also being quickly subducted, and Oceans iii and iv are following them. But Ocean ii’s subduction on the other side of the world helps keep their mid-ocean ridges balanced in the center.

To the south, 3 continues to collide with G over the south pole, and subduction of the ii-iii mid-ocean ridge causes a Laramide-type orogeny on its north coast.

Again like last time, at this point we should expect high global sea levels and inland seas—though perhaps that may be moderated by the presence of a large, likely-glaciated mountain range over the south pole.

50 mya: Late Andrewsian

DJ rifts from BFHI at 80 mya, and both this rift and the A-CE rift have several associated hotspots—as well as one final LIP in CE’s south.

G is moving pretty swiftly at this point, and as it shifts north from the south pole it creates something of a pocket between it and BFHI that Ocean ii is pushing into—such that there is a small section where the oceanic crust attached to BFHI and G is subducting under Ocean ii. The collision with Microcontinent 3 also continues, and has now enclosed a section of ocean, where slab rollback could lead to interesting patterns of land and sea.

Ocean vi has almost completely subducted away, and there is also a small section of Ocean vii remaining but it’s so small that I was too lazy to bother labelling it.

I did bother to add the vi-vii rift in this one, and I've noted an early rift pulling microcontinent 4 towards CE.

0 mya: Ostromian

At 30 mya, BF rifts from HI—the last rifting in the breakup of the supercontinent. At the same time, CE barrels into DJ, forming a broad Himalayan-type orogeny. We’ll use these events to mark the close of the Andrewsian and opening of our final period, the Ostromian.

Just before the collision, Microcontinent 4—from back in the Owenian—is pulled from DJ’s south coast and collides with the island arc to CE’s north, then is carried back north into DJ. This, in combination with slab rollback in the enclosed sea, should lead to an interesting, Mediterranean-like coastline.

Microcontinent 3 continues to collide with G, but by 0 mya there is still a substantial enclosed sea remaining.

Ocean ii continues to push into the pocket between G and HI. It isn’t a separate plate yet, but it will be before long so we can go ahead and call it Ocean viii. There’s also a small new section of subduction between A and G due to G’s counterclockwise rotation.

And because this is the last section of tectonic history, I’ve also pulled out some sections of the east coast of DJ, as it’s been moving slowly and we can expect some back-arc extension.

And that finally brings us to the “present day”. Here's a chart I threw together to help us keep track of the periods I've established and drift of the continents throughout them:

Widths are not proportional to continental area because I only realized I should have done that after I finished.

And indeed we can pretty well see a supercontinent cycle of breakup (Cuvieric, Anningic), assembly (Owenian, Huxleyic), Tenure (Marshian), and breakup once more (Copian, Andrewsian, Ostromian).

Now, we can’t really go on calling our continents “HI” and “BF”, so I’ve come up with some interim names for our 6 continents (I’m still counting CE and DJ as separate, given the high mountain range dividing them) and 4 major oceans until we figure out what sort of names their eventual inhabitants might use.

As a final touch for this post, we can use our recorded history to get a rough sense of the modern topography. As we’ve moved along, I’ve been marking out the areas affected by orogenies and LIPs, and now we have a record for all the major mountain-building events of the last 850 million years (You can use gplates to generate these age maps for you):

This map is in an equirectangular projection, which will be more useful for modifying it later on.

Mountain range height doesn’t directly correlate to age, but there is a relationship; Based on the heights and ages of many of Earth’s current ranges, average mountain height tends to begin around 2500 meters and drop about 5 m every million years, with peak height starting around 4000 m and dropping 8 m/myr. But there are plenty of exceptions and local variation and initial height is particularly broad (and really an exponential decay might be a better model, but we're working in generalities here). With this as a guide and a bit of artistic interpretation, I’ve put together a rough pass of Teacup Ae’s global topography:

These are average elevations, mind; we'll add plenty of peaks and valleys later on.

I’ll talk a bit more in a later post about the exact considerations regarding different mountain types, erosion, and uplift, but for now this gives us a general idea of the planet’s terrain and will help in working out the climate.

And that about does it. There’s still a lot of cleaning up and fleshing out to be done: In addition to further work on internal topography, the coastlines need to be refined, the island arcs expanded, and other islands added. Bear in mind that the boundaries shown aren’t actually coastlines, but boundaries between continental and oceanic crust (roughly). Even when sea level is low, much of the margins of the continental crust sections will be submerged continental shelves—more on passive margins. Though, in this specific case, I’ve tallied up the areas of crust at the start and end of the sequence and found that continental area didn’t increase as I was aiming for (I should have added more land during island arc collisions). It actually decreased, though only by about 0.2%.  So I might expand out these shelves a bit so I can retain more land area.

But for now, it’s a pretty good result.

Orogenies are black when active then fade to white as they age, LIPs are orange areas, hotspots are purple dots. Some of the borders can be a bit janky during collisions because I only bothered updating them every 50 million years.

An alternate view showing tectonic boundaries (red for divergent, blue for convergent, green for transverse) and distinct plates in different colors.

I’ll admit, this is a pretty tedious method to get a world map; what you see here is the result of a solid week of long afternoons (though it probably would have gone a lot faster if I was less obsessive about matching up all the lines just right). But I’m pretty happy with the result: A world like Earth, with a similar diversity of landforms, and a solid geological history to work off of (we’ll see the utility of that over the next few sections) but different in interesting ways: 

• Except between Hutton and Lyell, there are no land connections between the continents as there are between most continents on Earth. Hutton and Steno might still have a land bridge in the depths of an ice age, but at such points it would be covered with ice sheets anyway. There are only thin seas separating Steno, Holmes, and Hutton, and (once I’ve expanded the islands) Lyell and Wegener, so there will be some exchange of plant and animal life and intelligent Teacupites. Still, natural and social divergence of the continents will be easier than on Earth—especially on Agassiz, which is completely isolated from the other continents.
• Hutton has a European-like west coast and Asian-like east coast, but a much thinner (though mountainous and probably arid) interior than Eurasia, which could bring these diverse regions into closer contact than in our world.
• Steno and Agassiz both have sections near the poles that we should probably expect to develop ice caps, but also extend into temperate climate zones. We don’t have that full gradient on any one landmass on Earth today.
• Holmes is an “island continent” somewhat smaller than Australia, but less arid and closer to other landmasses and likely centers of civilization.

Next post, we’ll take a quick detour into looking into alternate modes of tectonics that we might see on more alien worlds, and then we’ll move on to planetary climates and biomes.

In Summary

• Convection of heat in Earth’s interior drives plate tectonics, but the motion of particular plates is determined by patterns of subduction.
• Plates spread apart at divergent boundaries—rifts and mid-ocean ridges—come together at convergent boundaries—subduction zones and collisional mountain ranges—and slide past each other at transverse boundaries.
• Ocean crust subducts and is completely replaced on ~250 million year timescales
• Continental crust is formed at subduction zones and never subducts, though it can be deformed and compressed by collisions.
• Earth undergoes a cycle of supercontinent formation, tenure, breakup, and reformation in a new form roughly every 750 million years.
• Supercontinet cycles can be either extroversions, with continents spreading and consuming the exterior ocean, or introversions, with continents reversing direction and consuming the interior ocean—or a combination of both.
• The position of subduction zones is the primary determining factor in the motion of plates, though the influence of mantle plumes plays a role as well.
• Mountains form at orogenies due to subduction zone volcanism or continental collisions, and then gradually erode down afterwards.
• Regular leading-edge, slab-rollback trailing-edge, and flat-slab subduction zones produce different types of active margins.
• Broad, Himalayas-like plateaus are formed by repeated impacts of small continental fragments followed by a large, high-speed continent collision.
• Complex, Mediterranean-like seas are formed by trapping of oceanic crust between converging continents.
• LIPs, hotspots, and failed rifts are other tectonic features that can affect global geography.

Notes

So, I'm going to be opening a patreon for the blog. Though I enjoy doing all this, it’s a big time commitment and I could use a little support as my bills start looming. But I won’t be pushing it past adding some links to the front page and other posts. I don’t want anyone to think I’m relying on this, and for now I won’t be putting anything behind a paywall (and I don’t intend to ever put anything behind one permanently). I’m just testing the waters a bit and hoping for a little extra motivation. 

Scotese has assembled a list of plate tectonics “rules” that he uses in his paleogeography reconstructions—I tried to contextualize them and translate them all into more layman’s terms throughout this post, but the originals are worth a read.

If you want to see a model of Earth's development from the very beginning, this video is a pretty decent representation, but with the major caveats that any geographic or atmospheric data before about 2 billion years ago is very speculative, and the landmasses shown don't necessarily represent all the landmass that existed at the time; just what we have direct evidence for.

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Part Vb