Tools & Methods
Here's a closer look at some of the tools and methods we're using while underway, courtesy of Jonathan Sleeper, a graduate student at the University of Hawaii. See also Chris Horvath's log entry on the day-to-day use of these instruments.
EM-122 Multibeam Sonar System
The primary mapping instrument on our current cruise is the Simrad EM-122 12 kHz multibeam sonar system. The multibeam sonar system consists of a “transmit array,” a linear array of transducers aligned parallel to the length of the ship that generate the sound-waves, and a “receiver array,” another array of transducers aligned perpendicular to the length of the ship that receive the reflected signal, together forming a “T” shape. A “multibeam” sonar system creates a fan-shaped pattern of sound waves that allows us to map a wide swath of seafloor, rather than just a single line directly below the ship as with the old single beam systems. Of course, basic physics tells us that sound travels as spherical waves, not linear beams. The “beams” are created where the adjacent spherical waves from each transducer constructively interfere with each other, concentrating the sound energy along linear rays, with gaps in between where they destructively interfere. For a more detailed explanation of the mathematics behind this beam-forming process and the design of sonar systems, this is one very useful reference: http://www.ldeo.columbia.edu/res/pi/MB-System/sonarfunction/SeaBeamMultibeamTheoryOperation.pdf. The EM-122 system forms 432 beams per ping across the swath, giving an approximate footprint (resolution) of 20 m at 2500 m water depth. The maximum angular swath width is 150°, although we typically used 130-140° during our survey. The width of this swath on the seafloor varies from 2.5-3.5x water depth (~3-10 km for our survey area), depending on angular swath width and seafloor geometry. A larger angular width covers a wider swath of seafloor, but also increases beam spacing and attenuation of the outer beams.
Watching the multibeam data come in! The top panel shows the swath of seafloor, with colors
denoting depth. The bottom panel shows a small patch of seafloor data looking along the ship track.
There are two primary types of data that can be recovered from the sonar system: bathymetry and backscatter. Sonar reflections in the water column are also recorded, but these are typically not useful for geophysical surveys or seafloor mapping. Bathymetry is simply a measure of the time it takes for the sound to travel to the seafloor and back to the receiver array. Knowing the speed of sound in water and how it changes with depth, the travel times can be converted into a distance measurement, and you can get an image of the topography (bathymetry) of the seafloor. Along with travel time, the receivers measure the intensity of the reflected sound waves (or echoes), which is also called backscatter. If the sound reflects off of a hard, smooth surface (a lava flow, fault scarp, or a sunken ship for example) the amplitude of the return signal is very high, but if the sound reflects off of a soft or rough surface (sediment for example), much of the energy is scattered and the intensity is weaker. The accuracy of both travel time measurements and amplitude measurements are also affected by their angular position within the swath. Beams toward swath edges typically have lower amplitude and are more subject to errors and noise, due to the longer travel distance and increased attenuation in the water column. Another factor that affects the amplitude of sonar returns is the seafloor geometry. If the seafloor is sloping away from the sonar beams, more of the energy is reflected away from the receivers and the returns are generally weaker. If the seafloor is perpendicular to the sonar beams, more of the energy is reflected straight back to the receivers and the return is stronger. This does not have a large effect on hull-mounted multibeam systems in deep water, but becomes a much more significant issue in shallow water or for deep-towed sonars that are towed within a few 100 m of the seafloor. Backscatter images essentially look like a black and white image of the seafloor and allow us to see structures such as faults and distinguish between fresh volcanism and sediments.
A sample screenshot of a backscatter image. The dark areas are more reflective (higher intensity).
Sonar data processing is generally a simpler process than magnetic and gravity processing. After extracting the data from the raw files, the first step is typically to identify and eliminate obviously bad pings (called ping editing). Sometimes one or more of the sonar beams will reflect off of something in the water column, or if the seas are rough it may lose track of where the real bottom is, resulting in depth values that can be very different than the surrounding values. These are usually easy to identify and eliminate and there are both automated and manual methods to do so. After the obviously bad pings have been eliminated, the data is interpolated onto a grid of x,y,z (longitude, latitude, depth) values, or for backscatter, depth is replaced with amplitude. The grid spacing is manually chosen based on the resolution of the system, the data point density, the water depth, and the amount of noise in the data, among other factors. Then a filter can be applied to both fill small gaps between pings (which become more prevalent in rough seas) and to further smooth out smaller amplitude noise in the data that was not eliminated during ping editing. To display the data properly, a color scale is chosen for the values, using colors for bathymetry and grayscale for backscatter, and this color scale can be customized depending on the range and distribution of z values in the data. For bathymetry it is also common to use illumination, where an angle is chosen for an imaginary “sun” in order to place highlights and shadows on the map and help the features to visually stand out.
Bell Aerospace BGM-3 Gravimeter
Gravimeters are used to measure the local gravitational field of the Earth. There are two main classes of gravimeters, those that measure absolute gravity and those that measure relative changes in gravity, which is what the BGM-3 gravimeter does. Typical absolute gravimeters use an accelerometer to measure the acceleration of a rising and falling mass in a vacuum. While some types have been constructed in a portable fashion to be used in field surveys, most absolute gravimeters are mounted in a fixed location and can be used to calibrate relative gravimeters. Relative gravimeters are much more commonly used for marine geophysical surveys and other practical applications. The simplest types of relative gravimeters are spring-based. They use a weight on a spring and carefully measure the amount that the spring stretches in response to local changes in gravity. Spring-based gravimeters require calibration to absolute gravity so we know the value of gravity when the spring is in its “resting” state. It also requires a variable conversion factor because the spring’s resistance increases as it is stretched further. The most accurate (and precise) relative gravimeter uses a superconducting Niobium sphere suspended in a stable magnetic field. As the gravitational pull varies on the sphere, the electric current required to keep it stable varies as well. The current is proportional to the Earth’s gravitational field and therefore can be used to calculate relative gravity. The BGM-3 uses a similar principal to the Niobium sphere gravimeter. The accelerometer is a “proof mass” wrapped in a coil of wire that is suspended between two permanent magnets and constrained so it can only move vertically. Changes in the position of the proof mass are detected and a current is automatically applied to stabilize the mass. The applied current is proportional to vertical acceleration and thus can be used to calculate the gravitational force acting on the ship. The gravimeter is mounted inside a temperature-controlled housing to eliminate errors due to temperature changes, and the instrument is typically mounted as close as possible to the center of the ship in order to minimize accelerations due to ship motion. The above information and further details on gravimeter design and the BGM-3’s performance during a marine geophysical survey are contained in the following reference: Bell, R.E., and A.B. Watts (1986), Evaluation of the BGM-3 sea gravity meter system onboard R/V Conrad, Geophysics, 51, No. 7, p. 1480-1493.
The Earth’s Gravitational Field
In basic physics classes, we typically think of gravity as a constant ~9.8 m/s2, but this is actually an over-simplification. In reality, gravity varies on a number of scales and has multiple components. At the global scale, gravity varies based on latitude because the Earth is not a perfect sphere. The Earth is actually an “oblate spheroid” that bulges slightly around the equator, meaning that the Earth’s radius is ~22 km larger at the equator than at the poles on average. Therefore, the equator is farther from the center of mass and thus gravity is slightly lower at the equator compared to the poles. Adding to this effect is the centrifugal force due to the Earth’s rotation. The equatorial region is rotating at a faster velocity than the polar regions, meaning that the centrifugal force, which pulls objects away from the Earth’s surface, also acts to slightly reduce gravity at the equator. These two effects combine to cause sea-level variation in gravitational acceleration to vary from ~9.780 m/s2 at the equator to ~9.832 m/s2 at the poles, meaning an object will weigh ~0.5% more at the poles. This variation in gravity is approximated by a “reference ellipsoid,” but this model does not account for smaller scale variations.
One source of smaller scale variations is tidal accelerations caused by the gravity of the moon, and to a lesser extent the sun. Tidal effects are cyclical and predictable, and therefore are possible to remove from the gravity measurements recorded by the gravimeter. Another source of smaller scale variations in gravity are density variations in the Earth’s crust and mantle. When performing a geophysical survey, these are the variations that we are interested in, because they reveal variations in subsurface crustal and upper mantle structure that can in turn shed light on the underlying processes that control them. During marine gravity surveys, the dominant short-period signal in the data is caused by the acceleration of the ship as it moves up and down on the ocean swells. These accelerations can be of the same order as the acceleration due to Earth’s gravity, and thus dominate the raw gravity signal. However, over the course of a few minutes, the ship’s motion averages to ~zero, so this effect is relatively easy to remove by averaging the values over a few minutes (we applied a 15 minute filter to the data for the current cruise). Another effect that must be accounted for if gravity measurements are made from a moving platform such as a ship or an airplane is called the Eotvos effect. As the ship moves eastward, it is moving with the rotation of the Earth and effectively increasing the centrifugal force, pulling the ship slightly away from the Earth and reducing the gravitational pull on the ship. The effect is opposite when moving westward. In order to remove this effect, we must determine the eastward component of the ship’s velocity and then use an equation to calculate the associated reduction (or increase) in gravity. One other effect that must be accounted for is the reduction in gravity due to elevation, called the free-air correction. For marine surveys, this is typically very small, since the gravimeter is mounted within a meter or two of the ocean surface. However, in terrestrial or aerial gravity surveys this can have a significant effect. An increase in elevation from sea level to 9,000 m (close to the top of Mt. Everest) will cause a weight decrease of ~0.29%.
After applying a filter to average out ship motions, removing the reference ellipsoid, applying the Eotvos correction, and applying the free-air correction, the result is called the “free-air anomaly.” The Bouguer correction is another common step in gravity processing. On land, this correction accounts for the mass of earth materials between the measurement location and the reference ellipsoid (roughly sea level), which is ignored if just the free-air correction is applied. The simple Bouguer correction calculates the effect on gravity from an infinite flat slab of a specified density, while the complete Bouguer correction takes into account local variations in topography. In a marine survey, this correction amounts to replacing the ocean with a slab of rock of a specified density (usually 2.8-3.0 g/cm3 for basalt). When this correction is applied, the resultant anomaly is the Bouguer anomaly, which allows us to see how the density structure varies relative to a typical density value. Lastly, prior to and following a gravity survey, the relative gravimeter used onboard must be calibrated to absolute gravity. This is done by a “gravity tie” to a location on land near the harbor, where the absolute gravity value has been measured precisely. A measurement is taken with a portable gravimeter directly next to the ship while in port and at the location where absolute gravity is known, and then this offset is applied to the known value of absolute gravity to get the absolute gravity value next to the ship. This value is then used to calculate the “DC shift” which can be added to the relative gravity values from the gravimeter to give values of absolute gravity. This is done both at the beginning and the end of the cruise in the exact same locations to see if the shipboard gravimeter values have drifted at all during the course of the survey. If so, the drift is assumed to be linear, and a drift correction is also applied to the gravity values.
Geometrics G-882 Cesium Vapor Magnetometer
The G-882 is a particularly sensitive type of optically pumped scalar magnetometer. A scalar magnetometer just measures the total strength of the magnetic field, while a vector magnetometer measures both the strength and the orientation of the field relative to the instrument. A Cesium vapor magnetometer works by passing photons through (hence the “optically pumped” part of the name) Cesium vapor in a sealed chamber. Electrons in the Cesium atoms with the proper spin orientation are excited to a higher energy level (orbital). The electrons then emit a proton and fall back to a lower energy state, but may or may not have the same spin. If the electron returns to the original spin state, the process repeats, but if it changes to the opposite spin, it can no longer absorb the photons and they all get through to the detector on the opposite side of the chamber. Eventually (in a matter of milliseconds) all of the electrons align in the opposite spin state and all photons pass through (i.e., the atoms become “polarized”). Then a small AC magnetic field of a certain frequency (called the Larmor frequency) is applied to change the state of the electrons and allow them to once again absorb photons. This frequency is precisely measured and since it is proportional to the external magnetic field, the external field can be calculated with a simple conversion factor. The instrument makes measurements at 10 Hz (10 measurements per second), which is much faster than we really need, but provides excellent data density to aid in filtering and cleaning up the data. The magnetometer is towed ~200 m behind the ship to avoid interference from ferrous materials in the ship.
Our magnetometer, otherwise known as "Maggie," getting a quick inspection.
The Earth’s Magnetic Field
The largest component of the total magnetic field is due to the Earth’s dipole field, produced by the flow of metallic elements (mostly iron and also nickel) in the liquid outer core, although this field is not a perfect dipole and changes in intensity and polarization over timescales of years to millions of years. Changes at this scale are modeled in the International Geomagnetic Reference Field (IGRF), which is updated every 5 years. This modeled field is removed from the total field measurements during processing of the data in order to isolate the smaller components of the magnetic field. Smaller-scale perturbations on scales of seconds to weeks can be caused by ionospheric effects or solar magnetic storms, which are much more difficult to model and therefore more difficult to remove from the total field measurements. The smallest component of the field and the piece that we are attempting to measure is the crustal magnetic field, formed by magnetic minerals (particularly magnetite) in seafloor rocks (typically basalt). While the rock is still molten, these minerals align with the Earth’s magnetic field, and when they cool below a certain temperature (called the Curie temperature), they remain frozen in the orientation of the Earth’s magnetic field at that particular time and create their own magnetic field in a process called thermoremanent magnetization. The intensity of the crustal field is largely governed by the magnetite content of the lavas as well as the strength of the Earth’s field, but as the lavas age the intensity decreases due to weathering, meaning that fresh lavas tend to have the highest magnetization intensities.
The other piece of information recorded by these minerals is the orientation of the magnetic field. The Earth’s magnetic field has switched polarity hundreds of times over its history, and the rocks formed at various points in Earth history still record the orientation of the field at the time they formed. Not only does it switch polarity over time, over shorter timescales the magnetic poles “wander” spatially in relation to the geographic poles. The angular difference between the geographic north pole and the magnetic north pole is called the declination. The inclination (also called the magnetic dip) is a measure of the angle of the magnetic field relative to the surface of the Earth. It varies with latitude from -90° (pointing straight up) at the magnetic south pole, to 0° (horizontal) at the magnetic equator, to +90° (pointing straight down) at the magnetic north pole. A scalar magnetometer only measures the intensity of the field, so it cannot provide information on declination and inclination, but clearly shows when the polarity reversals occur (the values change from positive to negative). Along spreading centers, these reversals create stripes of positively or negatively magnetized rocks as they form at the spreading center and migrate further off-axis. The observation and interpretation of these “stripes” was one of the most important pieces of evidence that established the existence of seafloor spreading and helped to confirm the theory of plate tectonics. By dating lava flows near these reversal boundaries and measuring how wide the strip of crust is between them, we can determine both current and past spreading rates. For instance, the last reversal has been dated at 780,000 years ago, so if we know that the reversals on either side of the current spreading axis are separated by 78 km (to make the math easy), we can estimate the average spreading rate at 10 cm per year for the last 780,000 years.
Jonathan Sleeper is working on his PhD in Marine Geology/Geophysics at UH Manoa. His current research focuses on backarc basin extensional tectonics, using backarc spreading center morphology along with associated volcanic and tectonic structures to try to understand the complex processes in subduction zones.