Category Archives: Geophysics

Which satellite is that? What does it measure?


Space may well be the final frontier (there are one or two on earth that still require some work), but the space around our own planet is decidedly crowded. Folk at NASA’s Goddard Space Center (Maryland) estimate about 2300 satellites now orbit Earth; vehicles in various states of repair, use or disuse, of which a little more than 1400 are operational Continue reading


The earth moved; GPS, earthquakes, and slow-slip


It is often useful to know where you are, in a spatial sense. In the old days (LOL), field geologists, the kind that make maps of rocks and earth structures would, armed with topography maps and compass, determine their location from some vantage point using line-of-sight and triangulation.  I don’t hanker for a return to these days. I’m grateful for the kind of location data instantly available on my smart phone – the little blue dot that seems to follow my course across some digital representation of the universe. But I acknowledge a kind of smugness, in the event the digital world nosedives, knowing that I can still find my way; no General Panic Stations (GPS) if the satellite-based Global Positioning System (the other GPS) fails.

GPS devices can also be attached to bits of the earth’s crust.  This is useful because the crust, whether continent, sea floor, or volcanic island, is always on the move. Continue reading


So, adding CO2 does increase surface heating; how science has filled another gap in our knowledge


Read any scientific paper or blog on climate and you’re bound to come across the phrase radiative forcing.  Radiative forcing is central to all climate science. Radiation from the sun heats our atmosphere and earth surface.  Some of this radiation is reflected back to space. If there is a balance between incoming and outgoing radiation then average global atmospheric temperatures neither increase or decrease. However, if the balance is perturbed, climate will warm or cool. Radiative forcing causes climate imbalances.  Thus, volcanic aerosols tend to cool things off, decreasing albedo will tend to warm them. Continue reading


Volcanism does not cause glaciations; let’s turn this statement on its head


It is almost a truism that volcanic eruptions influence climate. Cold winters and even failed crops, particularly in the northern hemisphere, followed on the heels the Tambora, Krakatoa, and Pinatubo eruptions.  But these climate aberrations were relatively short-lived, counted in years; the stratospheric aerosols and fine volcanic ash that reflect solar radiation back into space, eventually succumb to gravity and fall to earth.  Eruptions of this kind do not result in long-lived, or permanent changes; they are temporary blips on an evolving earth and an evolving climate. Continue reading


Tsunamis behave as shallow-water waves


Tsunami statistics make grim reading, which is why I am not going to quote any.  There are some great documentaries and websites that will regale you with all the stats you need. There’s even a couple of movies, where, if you sift through the hype, you may see a smidgen of science, or hear a bit of terminology added to the dialogue to give the impression of knowledgeable heroes.

The word Tsunami derives from two Japanese words; Tsu meaning harbour, and nami wave; an appropriate etymology given that these forces of nature really come into their own along shallow coasts and harbours. About 80% of tsunamis are generated by powerful earthquakes (particularly those beneath the sea floor); the remaining 20% result from large landslides, volcanic eruptions, and less frequently (fortunately) meteorite impacts. They are sometimes referred to, incorrectly, as tidal waves. Tides result from astronomical forces.  We can think of the succession of high and low tides as the passing of a wave that has a period of about 12 hours (the time from one high tide to the next). Tidal waves move along coasts such that a high tide at one location (i.e. the crest of the wave) will occur at a different time to that at a more distant location.  Tides also move water masses; waves do not.

Sea and lake surface waves are generated by wind. The wind provides the energy which is transferred to surface waters.  As a general rule, the stronger the wind, the greater are wave amplitude, wavelength, and speed. Water particles beneath waves have a circular or elliptical motion (referred to as orbitals); the larger circles occurring immediately below the crest, and decreasing in size to a depth that equates to about half the wavelength.  This means that in deep water, waves do not interact with the sea floor. This kind of surface wave is given the name deep-water wave, the speed of which depends only on the ratio of wavelength to wave period. Deep-water waves occur where water depth is greater than half the wavelength.

As waves approach the coast, the wave orbitals begin to touch the sea floor (also referred to as wave-base) and wave speed decreases.  At these depths (depth is less than half the wavelength), loose sediment can be moved by the wave orbitals. Some energy is transferred to the sea floor, but to conserve energy, the height, or wave amplitude must also increase. As you can see in the diagram, the orbitals also become flattened. At this stage, the waves have become shallow-water waves.

Although it may seem counterintuitive, tsunamis behave as shallow-water waves. They have long wavelengths, commonly measured in 10s to 100s of kilometres. The speed of shallow-water waves, including tsunamis, is independent of their wavelength, but is dependent on water depth in the following way:

Speed = (g . depth) (g = gravitational constant, 9.8m/s2; depth in metres)

In the case of tsunamis, the wavelength is many times greater than water depth, even in oceans more than 4000m deep. For example, a tsunami traveling across ocean that is 4000m deep will have a speed of 198m/second, or 713 km/hour. This animation of the 2010, M8.8 Chile earthquake and tsunami gives an impression of the speed of wave propagation across oceans, and the shape of the wave fronts. Tsunami waves commonly pass unnoticed beneath ships at sea or offshore rigs.  As they approach shallower water, their speed decreases to between 40-80km/hour (because speed is dependent on water depth), but the amount of energy in the wave changes very little; to compensate, the wave amplitude must increase. Earthquakes that generate tsunamis create several waves that spread out from the epicentre. All these waves can be destructive, and in some cases the first wave is the least harmful. It is also possible for a wave trough to reach the coast before the first wave crest; this results in a rapid drawdown of the water-level, exposing parts of the foreshore that would not normally be seen at even the lowest tides. Unfortunately, in all too short a time, the absence of water is replaced by a more menacing prospect.

Landslides can also produce monster waves; Lituya Bay in Alaska, 1958 is a good example with first-hand witnesses to the 15-22m wave. A prime example of volcanic eruption-derived waves is the cataclysmic 1883 Krakatoa eruption; a 30m tsunami wreaked havoc in Indonesia and across Sunda Strait.

Tsunami warning systems generally involve an international effort to, in the first instance, detect and pinpoint the epicentre of large earthquakes, and secondly, to detect tsunamis and predict their arrival times at different locations. There is a particular focus on submarine and near-coast, shallow crust seismic events of magnitude 7 and greater; high magnitude earthquakes deeper than about 100km generally do not produce destructive tsunamis.  Tsunami detection buoys have installed in 59 deep ocean locations, most around the Pacific rim.  The map shows the buoys to be located along tectonically active plate margins, such as the west coasts of North and South America, the Aleutian Arc, and other volcanic arcs – subduction zones from Japan through to New Zealand.

The deep-water buoys are anchored to the sea floor; for each sea-bottom buoy there is a linked surface buoy that relays data via satellite.  The deep buoys measure subtle changes in water pressure that can be used to calculate changes in sea-surface height.  The latest models have two-way communications so that a particular buoy can be programmed to search for pressure changes if an earthquake is known to have occurred.  Of course, all this is fine if a region has several hours to prepare for possible inundation.  Those close to epicentres may only have a few minutes to react.

The technology for tsunami prediction and warning is always improving. This is particularly the case for new generations of satellite that are tasked with collecting all manner of climate-related data, data relating to short- and long-term sea-level changes, and subtle changes in gravity and magnetic fields associated with earth’s ever-changing profile.

National Tsunami Warning Centre

Some Tsunami video clips

Boxing Day tsunami 2004 (Cornell Univ. animation)


There are two sides to every fault


In 1940-41, Harold Wellman, a creative but somewhat irreverent New Zealand geologist, along with his colleague Dick Willett, discovered a remarkably long, linear fault striking slightly oblique to, and a few kilometres landward of the South Island west coast; almost the entire length of the island. They called this massive structure the Alpine Fault. The Fault can be traced overland some 600 km, about 450km of this is a more-or-less single fault strand; at its northern extent the fault splits into several strands, all of which are active.

Most New Zealand geologists in the 1940s had little problem with a structure like this – admittedly it was very long, but most were familiar with faults, especially active ones.  By 1948 it had generally been accepted by the scientific community. The community did however have an issue with Wellman’s next discovery.  He realised that a certain group of rocks in the southern part of South Island (Otago region), were almost identical to a group at the north end of the Island (Nelson region).  He postulated in 1949 that these two geological domains were once a contiguous unit but had been separated some 500km by the Alpine Fault.  To many geologists at the time, this was going a bit too far, and it took several years to dispel the initial disbelief, and perhaps the odd conniption fit; one of the main criticisms was the absence of any reasonable mechanism to accomplish this geological feat. This was pre-Plate Tectonics, a time when many earth scientists still considered vertical movements of the earth’s crust to be the most important (although Alfred Wegener’s ideas on Continental Drift were discussed – it seems that Wellman was quite keen on this hypothesis). Fast forward to 1965 and a paper by J. Tuzo Wilson published in Nature, described a “New Class of Faults…”; Transform Faults.  Wellman’s discovery was about to acquire a mechanism, and become an iconic part of the new Plate Tectonics.

All plates identified by Plate Tectonic theory have boundaries, of which there are three basic types:

  • Spreading ridges and rifts, where upwelling magma creates new crust that moves away from the ridge,
  • Deep ocean trenches where two plates converge, forming a subduction zone that recycles old crust and mantle, and
  • Transform faults where two plates slide past one another. Most of this sliding is horizontal. If the movement between two massive slabs of crust and mantle were continuous then there would be few problems, other than a gradual (mm/year) change in one’s property boundary lines. But most movement along these fundamental structures is not continuous or uniform; it takes place in fits and starts – during earthquakes that commonly are very high magnitude, destructive events.

The Alpine Fault, and its close relative San Andreas Fault on the other side of the Pacific Ocean, are transform faults.  They each mark a boundary between two plates – if you walk across the San Andreas fault you pass from the Pacific Plate to the American Plate; over the Alpine Fault, from the Pacific to the Australian Plate.  There aren’t many places on earth where one can easily straddle two tectonic plates; these two transform faults provide great opportunities to become one with plate tectonics.

The Alpine Fault is geologically young.  The 500-km fault separation of the two geological domains began about 25 million years ago; from a geological perspective, this is really fast – for tectonic plates.  The west side of the fault moves northwards relative to the east side; it is referred to as a dextral (right-moving) strike-slip fault. At the same time, stresses acting against the fault have uplifted the landmass; over the last 12 million years, rocks formerly 20-30km deep, were pushed to the surface, forming the Southern Alps.  Coincidentally, erosion and glaciation have carved the landmass into the rugged mountain range that extends almost the full length of South Island. Averaged over the last 2 million years, the central part of the Alpine Fault has moved horizontally at a phenomenal 27mm/year, and vertically at 10mm/year.  It is thought that this extreme displacement of the earth’s crust is the result of large, M (magnitude)7.5 to M8 earthquakes occurring every 200-400 years, the most recent in 1717AD.

At its northern extent, the Alpine Fault splits into several large, active faults, some heading offshore, others into the southern North Island (the North Island Fault System) and these have been the focus of many destructive earthquakes in the M6 to M8 range.  More than 6m of horizontal displacement registered the M7.8 event along Kekerengu Fault in November 2016 (Kaikoura earthquake).  On January 23, 1855, up to 18m of horizontal displacement occurred during the Wairarapa Earthquake, estimated to have been M8.2 – M8.3.  The epicentre was only a few kilometres south of Wellington city, which suffered significant damage although few fatalities; there was also a tsunami that in places had a 10-11m run-up.

San Andreas Fault is another iconic example of earth’s major fractures, and probably the most intensely studied. It is about 1200km long, and like its New Zealand counterpart, consists of a master fault with many divergent, active and inactive fault strands. It began to move things around about 28 million years ago and has continued to do so ever since, coming to public prominence on April 18,1906 with the San Francisco M7.7 to M7.9 earthquake (and subsequent conflagration); one of the largest events along this fault.  Earthquake recurrence intervals vary along the San Andreas fault system; in the southern part it averages about 150 years, but in some fault segments like Big Bend, it may be as low as 100 years.

A commonly used method for estimating earthquake recurrence interval is to date young sediments that have accumulated close to faults. Silts and muds that accumulate in river or lake beds will frequently contain peats or fossil soils, layers of woody material, and sometimes volcanic ash; along coasts, beach deposits may be raised by successive earthquakes, and these too may contain shells, wood or bone.  These materials can be dated using carbon-14 and other dating techniques. The trick is to find layers that show some disturbance (for example from ground shaking, or displacement by actual faults) and then determine their age. There is always a fair degree of slop in recurrence numbers, a bit like predicting 1-in-500-year flood events (you might end up with 2 events in the space of 12 months!).

Serious earthquakes are a fact of life on transform faults; after all, what do you expect when 10-20 kilometre-thick slabs of rock slide past one another. Recurrence numbers for major events (greater than M6 or M7) may have annoying statistical variation, but they are based on sound science. The sensible lessons learned when someone else’s backyard is reduced to rubble, like – be prepared, or, let’s do more science – are all too quickly forgotten.  I guess it’s easier to point fingers after the fact, than to be on constant alert.

J Tuzo Wilson’s 1965 paper A new Class of Faults and their Bearing on Continental Drift was published in Nature, v.207, p.343-347.


It only takes a moment; the ups and downs of earthquakes


Seismic metaphors, or seismic as metaphor? Seismic, a word that geologists and geophysicists traditionally thought was reserved for their use, has been purloined by politicians and social scientists to describe momentous shifts in things like public attitudes and voter propensities.  Seismic is the anglicized derivative of the Greek verb seien, (shake) and the word seismos (earthquake).  Apparently, it first appeared in English language writing about the mid-19th century. Personally, I find it satisfying that the social milieu sees fit to apply the scientific word in such a useful, metaphorical sense.

The word seismic is also a nice descriptor of our restless, physical world, especially the bits we live on.  Most of the earth’s crust is under stress, some parts more than others.  Most stresses are generated by the movement and jostling of tectonic plates, particularly at boundaries where plates converge, collide, or slide past one another.  Different parts of the crust respond to stress by bending; this includes seemingly hard, immovable rock. If the stress is removed then the deformed rocks return to their original states; this is referred to as an elastic response. However, if the earth materials are bent too far or too fast, they will break. An interesting analogue for this process, and a historical one, is the collapse of a Tacoma suspension bridge in 1940.

Here, steel girders and tarmac bent and twisted under stress until the deformation reached a limit (called the elastic limit), at which point the bridge failed. When the earth’s crust fails, the seismic event, or earthquake, can be devastating. Earthquakes are caused by the sudden brittle failure of rock under stress; the failure takes place along a fault, across which land (or sea floor) is moved up, down, or sideways.  The rapid displacement of rock masses produces pulses of energy, or seismic waves.  It is these waves that do the damage.

There are two main kinds of seismic wave; body waves that propagate through the earth’s interior, and surface waves that move along the earth’s surface. Body waves include a primary, or P-wave, and a secondary or S-wave.  P-waves travel fastest; they are also called compressional waves because they tend to push and pull materials as they propagate. Slower S-waves, or shear waves, produce a side-ways motion. Shear waves are not transmitted through fluids such as sea water, or the molten interior of the earth. The cartoon below illustrates how the earth reacts to these two wave types.  Surface waves are the slowest to propagate but they are also the seismic pulses that do most of the damage during an earthquake. The animation shows the earth motion for one kind of surface wave; Rayleigh Waves. Rayleigh Waves produce a circular, or orbital motion of earth materials at the surface, a bit like particle motion beneath sea waves.


Each type of seismic wave is identified by the speed at which it moves, and the kind of movement that sediment and rock are subjected to.  These differences are expressed on seismograph recorders.  On a typical seismogram (below),  P-waves arrive first, followed by S-waves.   S-waves tend to have lower frequencies than P-waves (more spread out on the graph), but higher amplitude.  Surface waves usually have the highest amplitude.

Earthquake magnitude (M) reflects the severity of ground roll and shaking, and on seismograms corresponds to the amplitude of the signal (usually of surface waves).  M is expressed as a number (M1.8, M4.6, M7.8) up to a maximum of 10 (10 might be caused by a large meteorite impact – when there isn’t much left).  The numbers, and therefore the magnitude scale, are logarithmic, such that a magnitude of 4 (104) is 10 times smaller than M5 (105), and 100 times smaller than M6 (106).  We can also think of the changes in magnitude in terms of the energy released during a quake. It has been determined, (as an approximate empirical relationship) that for every unit increase in magnitude, there is a 27.5-fold increase in energy. The difference in energy is also logarithmic, such that an M8 event releases 571,914 times more energy than an M4 event (the magnitude is 10,000 times greater).  These kinds of numbers demand a degree of respect for the twists and turns our earth can throw at us.

Seismograms also help seismologists determine the Epicenter of an earthquake; the epicenter is the map location rather than the actual location, or focus, at depth. The calculation makes use of the fact that P-waves are faster than S-waves, so that the distance to an epicentre is based on the difference in arrival times for each type of wave.  Distance in this case is the radius of a circle centered on the seismograph location.  To triangulate the epicentre, at least two more seismographs in different locations are needed – in reality there can be hundreds of seismographs that allow the calculation.  If circles are drawn about each seismograph, each with a different radius, they will intersect at the epicentre.  Variations in the structure of the earth mean that seismic wave velocities can vary, so that the circles may not all intersect at a single point.  However, the large number of seismographs around the world means that location of the epicentre is usually accurate.

The November 2016 earthquake in Kaikoura, New Zealand was a M7.8 event.  We live about 500km north of the epicentre.  At 11pm we felt definite shaking (P-waves).  A few seconds later the shaking increased significantly.  At this point trees were swaying and water was slopping over the edge of the pool; this was due to ground roll from the slower surface waves.  Other than an adrenaline rush, there was no damage for us; farther south it was a very different story.  One spectacular outcome was the abrupt, lateral 6m shift along part of the New Zealand coastline; the same coast was also uplifted 1-2m.

I’ve always thought that New Zealand is a great place to witness geology in action.  But sometimes it can go a bit too far.

An excellent technical paper on the Kaikoura event by Ian Hamling (GNS Science) and a host of co-authors, has been published in Science, March 26, 2017.