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Mid-Ocean Ridge Seismicity
MID-OCEAN RIDGE SEISMICITY D. R. Bohnenstiehl, North Carolina State University, Raleigh, NC, USA R. P. Dziak, Oregon State University/National Oceanic and Atmospheric Administration, Hatfield Marine Science Center, Newport, OR, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The mid-ocean ridge that divides this planet’s ocean basins represents a rift system where new seafloor is emplaced, cooled, and deformed. These processes facilitate hydrothermal circulation and the exchange
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of elements between the solid Earth and ocean, support complex ecosystems in the absence of sunlight, and form one of the most active and longest belts of seismicity on the planet (Figure 1). This article outlines the tools used to study earthquakes in the oceanic ridge and transform environments, discusses the underlying mechanisms that cause or influence seismicity in these settings, and explores the impacts of earthquakes on submarine hydrothermal systems. As shown below, earthquakes can be used to track a number of important physical processes, including tectonic faulting, subsurface diking, seafloor eruptions, and hydrothermal cracking. As such, studies of earthquake patterns in space and time have become
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Figure 1 Global map of seismicity from National Earthquake Information Center (NEIC) catalog, MZ5, 1980–2005. Depths o33 km (red), 33–150 km (blue), and 4150 km (green). Mid-ocean ridges and oceanic transforms are defined by narrow bands of shallow hypocenter earthquakes. Spreading centers and approximate full spreading rates: Southern East Pacific Rise (SEPR, B140 mm yr 1), Northern East Pacific Rise (NEPR, B110 mm yr 1) Pacific–Antarctic Ridge (PAR, B65 mm yr 1), Galapagos Spreading Center (GSC, B45–60 mm yr 1), Chile Rise (ChR, B50 mm yr 1), Northern Mid-Atlantic Ridge (NMAR, 25 mm yr 1), Southern Mid-Atlantic Ridge (SMAR, B30 mm yr 1), Carlsberg Ridge (CaR, B30 mm yr 1), Central Indian Ridge (CIR, B35 mm yr 1), Southwest Indian Ridge (SWIR, B15 mm yr 1), Southeast Indian Ridge (SEIR, B70 mm yr 1), Kolbeinsey/Mohns Ridges (KR, B15–20 mm yr 1), Reykjanes Ridge (RR, B20 mm yr 1), Juan de Fuca and Gorda Ridges (JdFR/GR, B60 mm yr 1). Earthquake data from http://earthquake.usgs.gov.
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fundamental in efforts to understand the seafloor spreading system within an integrated framework.
Methods of Monitoring Seismicity Seismicity in the oceanic ridge-transform environment is monitored using a combination of hydroacoustic technologies, which record water-borne acoustic phases associated with submarine earthquakes, and traditional seismic sensors, which record ground motion induced by compressional and shear waves propagating through the solid Earth (Figure 2). These technologies should be viewed as complementary, as each presents certain advantages and limitations. Seismometers
There are hundreds of seismometers deployed on the continents and islands across the globe. These instruments, operated primarily by governments and universities, form networks of sensors that can be used to monitor seismicity, as well as clandestine nuclear tests. However, for spreading centers that lie thousands of kilometers from the nearest seismic station, our ability to detect and locate earthquakes remains limited. Within the most remote ocean basins, global seismic networks consistently detect only earthquakes larger than roughly magnitude (M) 4.5–5, or ruptures having an approximate physical
scale of Z1 km. The signals generated by smaller events typically attenuate to below background noise levels as they traverse long solid-Earth paths between the source region and land-based sensors. Seismometers may also be deployed on or beneath the seafloor in containers that protect the instrumentation from the extreme pressure of the deep ocean (Figure 2). Seafloor instruments are commonly known as ocean bottom seismometers (OBSs). They typically are deployed from a surface ship, allowed to free-fall into position, and later located using acoustic ranging techniques. Upon retrieval, an acoustic switch is remotely triggered, causing the instrument package to release a set of anchor weights and rise buoyantly to the surface in the vicinity of a waiting ship. The detection capabilities of an OBS array depend on the number and distribution of stations. Deployment of a dozen or more instruments for year-long or multiyear observations are becoming increasingly common with recent improvements in technology. OBSs deployed at very local scales, arrays with apertures of only a few kilometers, may be used to monitor small cracking events (often with Mo0) in the vicinity of hydrothermal systems. Deployments of larger-aperture arrays, tens of kilometers across, are typically used to study volcanic and tectonic processes within a spreading segment and may provide a record of many earthquakes that would otherwise go undetected by land-based seismometers.
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Hydrophones
Due to the combined dependence of ocean sound speed on pressure and temperature, much of the global ocean exhibits a low-velocity region known as the sound fixing and ranging (SOFAR) channel. Seismically generated acoustic energy may become trapped in the SOFAR, where it propagates laterally via a series of upward- and downward-turning refractions having little interaction with the seafloor or sea surface. The attenuation due to geometric spreading is cylindrical (R 1) for SOFAR-guided waves, making transmission significantly more efficient relative to solid-Earth phases that undergo spherical spreading loss (R 2). Water-borne, earthquake-generated acoustic signals were first observed on near-shore seismic stations. As these signals arrived after the faster propagating primary (P) and secondary (S) solidEarth phases, they were termed tertiary (T). Today T phases are monitored most commonly by hydrophones deployed directly within the SOFAR. The first systematic effort to use hydrophone data to produce a continuous catalog of mid-ocean ridge earthquakes began in the early 1990s when NOAA/ PMEL gained access to the US Navy’s Sound Surveillance System (SOSUS), a permanent network of bottom-mounted hydrophone arrays within the Northeast Pacific. This enabled the real-time detection of T waves generated by seafloor earthquakes and reduced the detection threshold for seismicity at the Juan de Fuca and Gorda Ridges by almost 2 orders of magnitude. The geometry of the SOSUS network also provided a better azimuthal distribution of stations than could be accomplished using land-based seismometers. The station geometry, combined with the existence of a well-defined velocity model of the oceans, also yielded significant improvements in location accuracy. Early successes using SOSUS facilitated the development of moored autonomous underwater hydrophones (AUHs) that could be used to monitor global ridge segments. In this design, the hydrophone sensor and instrument package are suspended within the SOFAR channel using a seafloor tether and foam flotation (Figure 2). These instruments have been deployed successfully along mid-ocean ridge spreading centers in the Atlantic and Pacific Oceans, as well as back-arc spreading systems of the Marianas (W. Pacific) and Bransfield Strait (Antarctica). A typical deployment consists of only six to seven instruments that monitor B201 along axis and provide a consistent record of earthquakes with M Z 2.53.0. Although regional hydrophone arrays can provide improvements in detection and location capabilities,
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relative to land-based seismic stations, it is not presently possible to extract the focal mechanism directly from the T waveform. Similarly, only relative depth estimates are acquired through measurement of the T wave rise time, defined as the time between the onset of the signal and its amplitude peak.
Global to Regional Tectonic Patterns A view of global seismic patterns (Figure 1) shows ocean basin earthquakes to be narrowly focused along the spreading axis. Such events have shallow hypocenters and focal mechanisms that dominantly indicate normal faulting on moderately dipping (B451) ridge-parallel structures. Shallow-hypocenter earthquakes also cluster tightly along the oceanic transforms that accommodate the differential motion between offset ridge segments, with the sense of motion along these conservative plate boundaries being dominantly strike-slip on subvertical structures. These patterns reflect a stress regime arising from the motion of the plates and the geometry of their boundaries (Figure 3). Transform and Segment Boundary Seismicity
In the late 1960s, focal mechanism studies of oceanic transform earthquakes provided key evidence in support of plate tectonic theory. Early physiographic maps of the oceans had outlined the mid-ocean ridge system and identified places where this submarine mountain range appeared to be offset laterally. One might infer from these observations that the two ridge segments were once aligned and that their offset represents the cumulative displacement along the connecting fault system. In contrast, the then emerging theory of plate tectonics required the Fracture zone (inactive)
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Transform Abyssal hill faulting Spreading axis Plate B Figure 3 Schematic showing the style of faulting and seismicity associated with an idealized spreading center and transform plate boundary. Double black line indicates the axis of the oceanic spreading center. Thin gray lines show normal faults along the rift flanks. Arrows indicate the direction of relative motion between the plates. Focal mechanisms shown with shaded regions indicating compressive first motions.
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opposite sense of motion along these structures in order to accommodate the process of seafloor spreading along two ridge segments that maintain a near-constant offset through time. As Sykes showed in 1967, the first-motion direction of P waves radiated during transform earthquakes clearly supports the latter model. Today the orientation and slip direction of earthquakes are determined routinely for global events larger than M 5–5.5. These data, combined with detailed morphological observations, provide a more complete picture of the structure and seismicity of oceanic transforms and higher-order segment boundaries. Many transforms are segmented, stepping in an en echelon fashion, with pull-apart basins and more commonly intra-transform spreading centers defining the segmentation. As an earthquake propagates, offsets of sufficient size may arrest the rupture, limiting its spatial extent. The maximum size of a transform earthquake is observed to be a function of slip rate, with the largest transform earthquakes occurring at slow spreading rates (o40 mm yr 1) and having sizes of roughly M 7.5. At spreading rates 4100 mm yr 1, the maximum observed magnitude drops to below M 6.5. This reflects the thermal structure of oceanic lithosphere, with a shallower depth of seismogenic rupture at faster spreading rates, and potentially a greater tendency for transform segmentation. Earthquakes with an oblique sense of slip and orientation are sometimes observed at inside corners, the region where the active transform segment meets the ridge axis (Figure 3). Such events may represent slip on curvilinear abyssal hill faults that rotate up to 15–201 as they approach the transform. In other cases, inside corner earthquakes may be better depicted as a compound rupture, where strike-slip motion along the transform and dip-slip motion along an orthogonal abyssal hill fault occur simultaneously. In these cases, the radiation pattern cannot be well described by the double-couple force model used to represent the vast majority of global earthquakes. Transform seismicity is restricted principally to the region between the two active spreading centers, as no relative motion is required across the fault beyond this region. Fracture-zone scars, however, may be traced across ocean basins and continue to represent zones of weakness. In some areas, intra-plate stresses are sufficient to reactivate these structures, infrequently generating large earthquakes. Oceanic transforms exhibit similar kinematics to continental transforms, such as the San Andreas Fault in the United States or the Anatolian Fault in
Turkey. Oceanic transforms, however, show one major difference – an inventory of earthquakes along oceanic transforms documents a dramatic ‘slip deficit’, with the majority of all transform motion occurring aseismically. A prediction of the moment release rate along a transform can be obtained as P Mo/t ¼ nmlw, where n is the relative plate velocity, m is the shear modulus of the rock, l is the length of the transform, w is the down-dip width of the rupture, and t is the time period considered. Mo is the static seismic moment, which for an individual earthquake is the product of fault rupture area (lw), mean slip, and m. The width w is generally taken to correspond to the B600–6501 isotherm, as constrained by slip inversion results and laboratorybased deformation studies (Figure 4). When the sum of the observed seismic moment on a transform is compared with the expected moment, their ratio (or seismic coupling coefficient, a) is typically B1 for continental transforms, but much less than 1 along oceanic transforms. The average a for all oceanic transforms is roughly 0.25, indicating that three-fourths of all transform motion occur aseismically. Studies investigating the velocity dependence of a yield conflicting results. Such assessments are complicated by the presence of significant inter-transform variability in a even at a given spreading rate, uncertainty in the depth of seismic faulting, and the somewhat limited timescale of the observations. The most recent and detailed studies, however, suggest little-to-no systematic relationship with slip rate, or perhaps slightly lower a at fastspreading rates. Aseismic fault motions on the oceanic transforms are likely taken up by slow, creeping events. These slow ruptures are inefficient at producing seismic waves and so are termed quiet or silent earthquakes. It recently has been suggested that silent earthquakes may trigger large seismic quakes in neighboring portions of the transform that are more prone to stick-slip behavior. In this view, some seismic events on the transforms may be viewed as aftershocks of these silent earthquakes. At higher-order segment boundaries, propagating rifts and overlapping spreading centers accommodate ridge segmentation through a mechanism known as bookshelf faulting (Figure 5). Here the seafloor between the overlapping rift tips is deformed by simple shear, resulting in the rotation of the initially ridge-parallel seafloor fabric. The style of deformation is reminiscent of a stack of books on a shelf tipping over, with slippage between the ‘books’ occurring along the preexisting abyssal fault systems. Focal mechanisms indicate a strike-slip sense of motion, with one set of nodal planes trending
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700°C Figure 4 (a) Map of the equatorial Mid-Atlantic Ridge showing the Romanche and Chain transform faults. The focal mechanisms obtained in the broad-band body-wave modeling are joined to their NEIC locations (red circles). (b) The centroid depths of the earthquakes are plotted as circles with symbol size proportional to magnitude. Solid symbols represent earthquakes with the bestresolved parameters, and open symbols those with fewer data or poorer fits. All the depths are accurate to within 3 km. Isotherms are calculated by using a half-space cooling model and averaging both sides of the transform. Slip inversions of the 1994 and 1995 Romanche earthquakes suggest that the former ruptured from near the surface to 20-km depth, and the latter from 10- to 25-km depth. This is consistent with the B6001 isotherm controlling the maximum depth of seismic slip on these transforms. The 1992 Chain earthquake is the only event with a centroid within the crust, but it was large enough to have ruptured into the upper mantle. Reprinted by permission from Macmillan Publishers Ltd: Nature (Abercrombie R and Ekstro¨m G (2001) Earthquake slip on oceanic transform faults. Nature 410: 74–77), copyright (2001).
parallel to an abyssal fabric that becomes progressively rotated within the deformation zone. This mode of deformation is most common at fast to intermediate spreading rates, with well-studied examples along the southern East Pacific Rise (4120 mm yr 1), Galapagos (B60 mm yr 1), and Lau Basin spreading centers (B100 mm yr 1). A spectacular slow-spreading example, however, can be observed within Iceland’s southern seismic zone. Spreading-center Earthquakes
When the oceanic ridge systems are monitored using global seismic stations (MZB4.5) or regional hydrophone arrays (MZB2.53) a sharp contrast in the frequency of mid-ocean ridge earthquakes is observed as function of spreading rate (Figures 1 and 6). Along fast-spreading (480 mm yr 1) axial highs, the ridge displays a dearth of small- to moderatemagnitude earthquakes. In contrast, such events are
abundant along the rift valleys of a slow-spreading (o40 mm yr 1) ridge system. These first-order patterns largely reflect the variable thickness of the brittle layer, which controls the predicted moment release, and seismic coupling coefficient of the rift-bounding normal fault systems. The predicted rate of seismic moment release can be estimated in a manner analogous to that described P for oceanic transforms: Mo/t ¼ nmlw/sin ycos y, where y is the fault dip, n is the full spreading rate, l is the length of the plate boundary, and w is the thickness of the seismogenic lithosphere. A comparison with the observed moment release at a slowspreading rift zone indicates that only 10–20% of the plate divergence is accommodated by seismic processes. This estimate agrees well with the amount of brittle strain accommodated by the normal fault populations of the abyssal plains. Hence, riftbounding normal fault systems at slow-spreading centers display a high-level of seismic coupling, with
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Figure 5 Central Lau Basin propagator with focal mechanics solutions. The Central Lau Basin Spreading Center is propagating to the south as spreading ceases along the northern tip of the Eastern Lau Spreading Center. Inset shows idealized case of bookshelf faulting, with progressive clockwise rotation of the abyssal fabric as the left-offset-propagating rift advances to the south. Focal mechanisms adapted from Wetzel et al. (1993) and the Global Centroid Moment Tensor Project (http://www.globalcmt.org). From Wetzel LR, Wiens DA, and Kleinrock MC (1993) Evidence from earthquakes for bookshelf faulting at large non-transform ridge offsets. Nature 362: 235–237.
the ratio of seismic fault slip to expected fault slip being close to 1. However, the emplacement of dikes along or near the rift axis takes up the majority (80–90%) of the plate separation. At fast-spreading ridges, commonly less than 1% of the predicted moment release is observed seismically along the rift. While most of the plate separation is accommodated by diking, as it was at slower rates, faulting studies indicated the fastspreading lithosphere undergoes an extension of B4–8%. The observed moment release is insufficient to account for this. Fast-spreading normal fault systems, therefore, appear to have low a and must accrue displacement during aseismic slip events or by
abundant microseismic activity too small to be detected by the hydrophones or global seismic stations. At intermediate spreading rates, the density of seismic events shows a first-order correspondence with the morphology of the ridge, with small- to moderate-magnitude earthquakes being abundant along rifted-spreading centers and comparatively rare at intermediate-rate axial highs (Figure 7). Intermediate-spreading-rate ridges can therefore assume both the morphologic and seismic characteristics of the fast- and slow-spreading end members. The maximum hypocentral depth of a rift zone earthquake decreases with increasing spreading rate, reflecting the temperature limits of brittle faulting
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and steeper thermal gradients at faster spreading rates (Figure 8(a)). OBS studies define additional intra- and inter-segment patterns. Within some slowspreading segments, the along-axis depth of microseismicity shallows near the center of a segment and deepens near its ends. This reflects the presence of a hotter lithosphere near segment center and is consistent with a three-dimensional (3-D) pattern of upwelling and melt focusing at slow-spreading rates. Studies on the Mid-Atlantic Ridge also show a positive correlation between the relief of the median valley and the maximum depth of microseismicity detected using OBSs (Figure 8(b)). This is consistent with stretching models that indicate greater fault offsets in regions of thicker brittle lithosphere.
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Figure 6 Mid-ocean ridge seismicity viewed from regional hydrophone arrays. (a) Mid-Atlantic Ridge (B25 mm yr 1), (b) East Pacific Rise (B110 mm yr 1) and Galapagos Spreading Center (B45–60 mm yr 1). Events located using four or more hydrophones are shown. M Z 2.5 3 earthquakes are consistently detected by these arrays. Yellow stars indicate position of hydrophones in each array. Red arrows mark firstorder (transform) offsets of the ridge axis. Data from http:// www.pmel.noaa.gov.
Prior to an eruption, mid-ocean ridge melts accumulate in crustal-level chambers. At fast-spreading rates axial magma lenses are ubiquitous, being found beneath 460% of the axis that has been surveyed using multichannel reflection techniques. At slowspreading rates, crustal-level melt is absent beneath much of the ridge axis and melt bodies may be ephemeral or localized beneath axial volcanoes. At intermediate rates, the distribution of melt can range between the two extremes, with some axial highs showing nearly continuous along-axis magma chambers reminiscent of fast-spreading centers. The accumulation of melt within the crustal chamber elevates its pressure relative to the surrounding host rock. This inflation deforms the surrounding crust and triggers earthquakes within the vicinity of the chamber. For magma to leave the chamber, the pressure (P) within must exceed the sum of the minimum confining stresses (s3) at the chamber boundary and the tensile strength (T) of the rock. The dike’s orientation will be orthogonal to the least compression stress direction and therefore it should be subvertical and aligned parallel to the axis of spreading. Although cracking in the vicinity of the dike tip creates many small earthquakes, most events of sufficient size to be detected on global seismic or regional hydrophone arrays are thought to occur on preexisting fault surfaces. In the region above the dike, a narrow zone of ridge-normal extension exists where seismicity is localized (Figure 9). A broader zone of extension exists beyond the along-axis tips of the intrusion, where seismicity will be triggered in front of a laterally propagating dike. As the dike passes, the lithosphere will be compressed at depth within the region adjacent to the dike.
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Figure 8 (a) Centroid depth of mid-ocean ridge earthquakes obtained from body waveform inversion of teleseismic arrivals vs. halfspreading rate. Maximum depth of seismic faulting is inferred to be less than twice the maximum centroid depth. (b) Relationship between cross-axis relief and maximum depth of seismicity. Maximum depth of seismicity was inferred from the focal depths of inner valley floor earthquakes, as recorded by OBS studies. (a) Reprinted by permission from Macmillan Publishers Ltd: Nature (Solomon SC, Huang PY, and Meinke L (1988) The seismic moment budget of slowly spreading ridges. Nature 334: 58–61), copyright (1988). (b) Reproduced from Barclay AH, Toomey DR, and Solomon SC (2001) Microearthquake characteristics and crustal VP/VS structure at the Mid-Atlantic Ridge, 351 N. Journal of Geophysical Research 106: 2017–2034 (doi:10.1029/2000JB900371), with permission from American Geophysical Union.
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Much of what we know about mid-ocean ridge dike injection comes from SOSUS observations of intermediate-rate-spreading centers in the NE Pacific. On 26 June 1993, soon after monitoring began, an intense episode of volcanic seismicity was recorded in real time along the CoAxial Segment of the Juan de Fuca Ridge (JdFR), marking the first observation of such an event (Figure 10). The seismic swarm started near 461 150 N and migrated northward during the next 40 h to B461 360 N, where majority of earthquake activity occurred during the following 3 weeks. Earthquakes propagated NNE at a velocity of 0.370.1 m s 1. The reservoir that acted as the dike’s source likely resided beneath, or to the south of, the initial swarm of earthquakes. The character and propagation velocity of this earthquake swarm were very similar to dike injections observed at Krafla and Kilauea Volcanoes. Seismicity went undetected by land-based seismic networks, suggesting earthquakes Mr4.0. Since the CoAxial activity, six additional magmatic episodes have been observed hydroacoustically on the Northeast Pacific spreading centers, and several distinctive characteristics have been identified: (1)
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the sequences are swarms, lacking a dominant event, with a temporal history that cannot be described by a power-law decay in event rate (i.e., Omori’s law); (2) the total number of earthquakes during swarms exceeds 10% (4350 events/week) of the variance of long-term background JdFR–Gorda Ridge seismicity; (3) swarms have several episodes of intense activity that reach 50–100 earthquakes/hour; (4) swarms last from several (45) days to several weeks; (5) earthquakes may migrate up to tens of kilometers alongaxis following the lateral injection of magma through the crust; and (6) swarms may be accompanied by continuous, broad-band energy (3–30 Hz) interpreted as ‘intrusion tremor’, resulting from magma breaking through the crust. Autonomous hydrophone recordings from the north-central Mid-Atlantic Ridge (B25 mm yr 1) indicate that diking events are less frequent, consistent with the lower rates of plate separation. During more than 5 years of monitoring, only one probable volcanogenic swarm was detected along B2500 km of ridge axis. This activity was within the Lucky Strike segment (371 N), somewhat outside of the AUH array. Earthquake locations could be determined for 147 hydrophone-detected events, with 33 of sufficient size to be located by land-based seismometers (M 3.6–5.0). In terms of total moment release, this earthquake sequence was one of largest on the Mid-Atlantic Ridge in the last three decades. The activity displayed a swarm-like temporal
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behavior and was accompanied by intrusion tremor; however, it was of short duration (o2 days) relative to previous episodes detected on the Northeast Pacific spreading centers. By far the largest episode of volcanogenic seismicity recorded on a mid-ocean ridge system occurred within the Arctic Ocean basin on the ultra-slowspreading (o5 mm yr 1) Gakkel Ridge in 1999. Seismic activity began in mid-January and continued vigorously for 3 months, with a reduced rate of activity continuing for 4 additional months (Figure 11). In total, 252 events were large enough (M44) to be recorded on global seismic networks (no hydroacoustic monitoring was in place). Subsequent geophysical studies of the Gakkel indicate the presence of a large, recently erupted flow and a volcanic peak directly in the area of seismic activity, with several large central volcanoes spectacled throughout an otherwise magma starved and heavily sedimented rift system. The duration of the 1999 activity suggests intrusion event(s) spanning several months and indicates a significant volume of magma beneath these central volcanoes. Because of their small magnitudes, earthquakes generated by diking events along fast-spreading ridge crests may be difficult to detect using hydroacoustic or global seismic techniques. Despite the greater rate of plate separation, during an 8-year period of hydroacoustic monitoring on the East Pacific Rise (101 S–101 N), only a handful of short-duration
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Figure 11 Seismovolcanic activity on the ultra-slow-spreading Gakkel Ridge near 851 N, 851 E. (a) Three-dimensional view from west (bottom) to east (top). The dark, reflective terrain is centered about a close-contoured high having a maximum vertical relief of 500 m. Red circles show the locations of epicenters, Jan.–Sep. 1999. (b) Histogram showing progression of the swarm through time, with each bar representing the number of events per day. Inset figure shows the cumulative number of events through time. There is a clear decrease in the rate of activity on 15 April (dot-dash line). Dashed line shows 6 May, when the USS Hawkbill passed over the area collecting the side scan imagery shown in (a). (a) Reprinted by permission from Macmillan Publishers Ltd: Nature (Edwards MH, Kurras GJ, Tolstoy M, Bohnenstieh DR, Coakley BJ, and Cochran JR (2001) Evidence of recent volcanic activity on the ultra-slow spreading Gakkel Ridge. Nature 409: 808–812), copyright (2001).
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(hours to days) swarms containing a few tens to hundreds of locatable earthquakes were detected. They were primarily located along intra-transform spreading centers and near segment ends, where a somewhat colder, thicker lithosphere might be expected to support larger ruptures. In January of 2006, a submarine diking event and seafloor eruption were captured for the first time by a network of OBSs. The eruption was located within the well-studied 91500 N region of the East Pacific Rise. As part of a multidisciplinary study, a network of up to 12 OBSs was deployed at this site in October 2003. Analysis has shown a gradual ramp-up in activity during a more than 2-year interval prior to the eruption, with activity culminating in an intense period of seismicity that lasted B6 h on 22 January 2006. Although many of the OBS instruments were destroyed by the eruption, the event was also recorded by regional AUHs, and the joint analysis of these data is expected to further illuminate the process of dike injection at a fast-spreading ridge. It should be emphasized that these somewhat limited observations may not be characteristic of all eruptive activity at mid-ocean ridges. Although ridge scientists often refer to a volcano-tectonic cycle, given a sufficiently long period of observation it is likely that mid-ocean ridge eruptions, like their continental equivalents, will display a power-law scaling with respect to size and duration, with smaller episodes occurring more frequently than larger ones. The scaling exponent and maximum eruption size, however, should be expected to vary as a function of spreading rate, reflecting variability in the dynamics of each system and the size of its crustal reservoir. Likewise, eruption frequency is likely to show a long-term clustering, rather than the periodic behavior sometimes alluded to.
field, strain rates from cooling are estimated to be on the order of 10 6 yr 1, about 3 orders of magnitude higher than tectonic strain rates. In 1995, a vent-scale OBS study lasting 105 days was conducted at 91 500 N on the East Pacific Rise, where a series of temperature probes also were deployed within the high-temperature vent systems (Figure 12). This study recorded a swarm of 162 microearthquakes during a period of about 3 h. Hypocenters defined a subvertical column at a depth of 0.7–1.1 km (Figure 12). Four days later, temperature sensors within a high temperature (Bio9) vent began to record increasing fluid temperature of B1 1C/day, with this trend continuing for 7 days. Vent temperature then returned to background levels during the next B120 days. Comparisons with vent temperatures during more than 3 years of monitoring showed this postseismic fluctuation to be the largest recorded, making a strong argument that the earthquakes had perturbed the hydrothermal system. It was suggested that fluids rapidly penetrated the new fractures and extracted heat from the fresh rock. A decade later, multiyear OBS monitoring, in conjunction with in situ temperature and chemical and biological studies, has been reestablished in the 91 500 N region. The results of this ongoing work, combined with similar vent-scale OBS studies conducted at intermediate- and slow-spreading vent sites, show microseismicity to be ubiquitous in hydrothermal areas. Moreover, they suggest a more complex relationship between seismic activity and vent hydrology than could have been envisioned based on the 1995 study, with temperature and flow responses exhibiting variable amplitude, sign, and phase, and many swarms producing no detectable perturbation at the vent sites.
Hydrothermal Seismicity
Tidal Triggering
Within the mid-ocean ridge hydrothermal regime, earthquake activity may be triggered in response to the removal of heat, which leads to thermal contraction and subsequent cracking, or via hydrofracture when trapped pockets of fluid are heated. Such events are typically small, with rupture diameters of 1 m to tens of meters, and can only be recorded using local OBS arrays deployed within the high-temperature vent fields. Within the shallow crust the predicted mode of failure is for purely tensional (mode I) or mixed-mode fracture, rather than the shear failure observed for the largermagnitude spreading-center earthquakes. Given estimates of the hydrothermal heat flux within a vent
Solar and lunar tidal forces exert short period stress variations that act on the Earth. Beneath the oceans and in coastal areas, stress changes influencing earthquake occurrence reflect the combined effect of the direct Earth tide and indirect loading of the Earth by the ocean tides. Together these changes are on the order of 103 104 Pa, much less than the average stress drop of an earthquake or the strength of crustal rocks. Many studies have examined the role of tides in triggering seismicity around the globe. With the exception of some terrestrial volcanic areas, earthquakes and tides are generally not correlated or weakly correlated during periods when tidal fluctuations are at their largest amplitude.
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Figure 12 (a) Ridge normal cross section of relocated microearthquakes associated with a 1995 swarm beneath the 91 500 N hydrothermal site on the East Pacific Rise. (b) Correlation between the microearthquake swarm and the fluid exit temperature at vent Bio9. Reproduced from Sohn RA, Hildebrand JA, and Webb SC (1999) A microearthquake survey of the high-temperature vent fields on the volcanically active East Pacific Rise (91 500 N). Journal of Geophysical Research 104: 25367–25378 (doi:10.1029/ 1999JB900263), with permission from American Geophysical Union.
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Julian days, 1995 Figure 13 (a) Fifteen-day time histogram of the earthquake count in 2-h bins with the ocean tide time series superimposed. Earthquakes are triggered at periods of low ocean tide, when the seafloor is unloading and stress within the crust is most extensional. Data collected during an ocean bottom seismic experiment near 481 N, 2311 E on the Endeavour Segment of the intermediate-rate JdFR. Reproduced from Wilcock W (2001) Tidal triggering of microearthquakes on the Juan de Fuca Ridge. Geophysical Research Letters 28: 3999–4002 (doi:10.1029/2001GL013370), with permission from American Geophysical Union.
Similarly, analysis of hydrophone-derived and global seismic catalogs does not show a robust correlation between tidal phase and the occurrence times of small- to moderate-size mid-ocean ridge earthquakes. However, microseismicity (Mo2) within axial hydrothermal systems commonly shows a strong correlation, with earthquakes occurring
during periods of peak extensional stress. In some locations, such as the JdFR in the Northeast Pacific Ocean, the amplitude of the ocean tides is large and these peak extensional periods correspond to times of low tide, when the seafloor is unloaded as the height of the overlying water column reaches a minimum (Figure 13). In other areas, such as the
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91 500 N region of the East Pacific Rise, tidal changes in the height of the sea surface are minimal and the Earth tide stresses have been shown to dominate. One explanation for these observations comes from laboratory studies of rock failure, which indicate that earthquake triggering in response to a periodic loading is dependent on the period of the oscillation relative to the time it takes a slip event to nucleate. Earthquake nucleation time, which is inversely proportional to stressing rate, is estimated to be on the order of a year or more for typical tectonic settings. Consequently, these systems are largely insensitive to tidal stress changes of semi-diurnal frequency. In volcano-hydrothermal systems, however, cooling- and magma-induced stress changes elevate the stressing rate by several orders of magnitude. Therefore, earthquake nucleation times become
comparable to tidal periods, and microseismicity in this setting becomes susceptible to tidal influence.
Impacts on Hydrothermal Systems Earthquakes may disturb the hydrology of marine hydrothermal systems through several mechanisms. As described previously, microearthquakes occurring beneath a vent field can alter the local permeability, creating new fluid pathways that allow the migration of cold water into the reaction zone or redirect the escape of buoyant, heated waters. Since the precipitation of minerals within the up- and downflow zones is predicted to reduce permeability over time, earthquakes events may be critical in maintaining the longevity of hydrothermal systems.
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Figure 14 Map showing acoustically derived locations of foreshocks (red dots), mainshock (yellow star), and aftershocks (white dots) of 1–7 Jun. 2000 earthquake sequence on the Western Blanco Transform. Error bars on SOSUS earthquake locations represent one standard error. Note difference in location relative to NEIC seismic epicenters (black squares), which are scattered across the plate interior. Locations of Vent 1 and Plume hydrothermal vent sites along southern JdFR are also shown. (b) Twenty-three-month temperature record at Plume hydrothermal site. Fluid at Plume vent decreased 4.9 1C during the 18 h immediately following mainshock. Three additional temperature drops appear to be associated with much smaller (MB4.0 4.2) earthquakes along the western Blanco (blue dots in (a)). (b) Eleven-month temperature record at Vent 1 hydrothermal site shows a possible delayed response to the June 2000, M 6.2 earthquake. After 7 months of steady readings (27273 1C), probe temperature decreased slightly prior to the event, but then dropped markedly by 18 1C begining 7 days after the event. Reproduced from Dziak RP, Chadwick WW, Jr., Fox CG, and Embley RW (2003) Hydrothermal temperature changes at the southern Juan de Fuca Ridge associated with MW 6.2 Blanco Transform earthquake. Geology 31: 119–122, with permission from the Geological Society of America.
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Earthquakes also induce static stress changes, with amplitudes that decay with distance from the earthquake source and become insignificant at ranges beyond 1–2 rupture lengths. These changes trigger secondary earthquakes (aftershocks) and dilate (or compress) the surrounding lithosphere. At more remote distances, up to several tens of rupture length from the epicenter, the transient stresses associated with passing seismic waves also may impact the hydrothermal system. In these cases, earthquakeinduced ground shaking is thought to jar open some hydrothermal conduits and seal others. Observations in the marine environment indicate that seismicity may create fluctuations in vent flow, temperature, and chemistry with variable characteristics. For example, drops in temperature at vent sites on the southern JdFR have been observed to correlate with earthquakes on the Western Blanco Transform at distances 430 km (Figure 14). The sensitivity, phase, and amplitude of these changes, however, vary significantly even between closely spaced sites (Figures 14(a) and 14(b)). Given the distances involved, temperature changes correlated with moderate-size (MB4.0) earthquakes on the Blanco Transform likely reflect the systems response to dynamic stress transients, while much larger earthquakes (M46) also induced significant static changes. Importantly, earthquake-induced changes within mid-ocean ridge hydrothermal systems can dramatically affect the biological communities that derive their energy from chemosynthetic processes. For example, the sub-seafloor microbial community is extremely sensitive to temperature, oxygen content, pH, and other environmental variables. It thrives in subsurface zones where the hot hydrothermal fluid mixes with entrained seawater. Changes in crustal fluid temperatures of only a few degrees, or a minor alteration of crustal permeability, can cause the prevailing microbial species to weaken and encourage new species to thrive and become dominant. Similarly, changes in the effluent thermal flux of hydrothermal vents can enhance heat output changing the size and temperature of the thermal boundary layer, a zone just above the axial floor that supports abundant and diverse macrofaunal communities. Still more dramatic impacts may be associated with the emplacement of magma at depth or the eruption of lava onto the seafloor. The heat supplied by these systems feeds massive bacteria blooms (floc) and may generate event megaplumes within the water column, huge volumes of hydrothermal fluid enriched in reduced chemicals that rise up to 1 km above the seafloor.
See also Acoustics, Deep Ocean. Hydrothermal Vent Fluids, Chemistry of. Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Seismic Structure. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Seamounts and Off-Ridge Volcanism.
Further Reading Abercrombie R and Ekstro¨m G (2001) Earthquake slip on oceanic transform faults. Nature 410: 74--77. Barclay AH, Toomey DR, and Solomon SC (2001) Microearthquake characteristics and crustal VP/VS structure at the Mid-Atlantic Ridge, 351 N. Journal of Geophysical Research 106: 2017--2034 (doi:10.1029/ 2000JB900371). Bird P, Kagan YY, and Jackson DD (2002) Plate tectonics and earthquake potential of spreading ridges and oceanic transform faults. In: Stein S and Freymueller JT (eds.) Geodynamics Series 30: Plate Boundary Zones, pp. 203--218. Washington, DC: American Geophysical Union. Boettcher MS and Jordan TH (2004) Earthquake scaling relations for mid-ocean ridge transform faults. Journal of Geophysical Research 109: B12302 (doi:10.1029/ 2004JB003110). Dziak RP, Chadwick WW, Jr., Fox CG, and Embley RW (2003) Hydrothermal temperature changes at the southern Juan de Fuca Ridge associated with MW 6.2 Blanco Transform earthquake. Geology 31: 119--122. Dziak RP, Fox CG, and Schreiner AE (1995) The June–July 1993 seismo-acoustic event at CoAxial Segment, Juan de Fuca Ridge: Evidence for a lateral dike injection. Geophysical Research Letters 22: 135--138. Edwards MH, Kurras GJ, Tolstoy M, Bohnenstieh DR, Coakley BJ, and Cochran JR (2001) Evidence of recent volcanic activity on the ultra-slow spreading Gakkel Ridge. Nature 409: 808--812. Fox CG, Matsumoto H, and Lau T-KA (2001) Monitoring Pacific Ocean seismicity from an autonomous hydrophone array. Journal of Geophysical Research 106: 4183--4206. Sohn RA, Hildebrand JA, and Webb SC (1999) A microearthquake survey of the high-temperature vent fields on the volcanically active East Pacific Rise (91 500 N). Journal of Geophysical Research 104: 25367--25378 (doi:10.1029/1999JB900263). Solomon SC, Huang PY, and Meinke L (1988) The seismic moment budget of slowly spreading ridges. Nature 334: 58--61. Sykes LR (1967) Mechanism of earthquakes and nature of faulting on the mid-oceanic ridges. Journal of Geophysical Research 72: 2131--2153. Tolstoy M, Cowen JP, Baker ET, et al. (2006) A seafloor spreading event captured by seismometers. Science 314: 1920--1922 (doi:10.1126/science.1137082).
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Wetzel LR, Wiens DA, and Kleinrock MC (1993) Evidence from earthquakes for bookshelf faulting at large nontransform ridge offsets. Nature 362: 235--237. Wilcock W (2001) Tidal triggering of microearthquakes on the Juan de Fuca Ridge. Geophysical Research Letters 28: 3999--4002 (doi:10.1029/2001GL013370).
Relevant Websites http://www.globalcmt.org – Global CMT Web Page.
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http://www.pmel.noaa.gov – National Oceanic and Atmospheric Administration’s Acoustic Monitoring Program. http://earthquake.usgs.gov – US Geological Survey Earthquake Hazards Program.
Introduction The mid-ocean ridge that divides this planet’s ocean basins represents a rift system where new seafloor is emplaced, cooled, and deformed. These processes facilitate hydrothermal circulation and the exchange
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of elements between the solid Earth and ocean, support complex ecosystems in the absence of sunlight, and form one of the most active and longest belts of seismicity on the planet (Figure 1). This article outlines the tools used to study earthquakes in the oceanic ridge and transform environments, discusses the underlying mechanisms that cause or influence seismicity in these settings, and explores the impacts of earthquakes on submarine hydrothermal systems. As shown below, earthquakes can be used to track a number of important physical processes, including tectonic faulting, subsurface diking, seafloor eruptions, and hydrothermal cracking. As such, studies of earthquake patterns in space and time have become
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Figure 1 Global map of seismicity from National Earthquake Information Center (NEIC) catalog, MZ5, 1980–2005. Depths o33 km (red), 33–150 km (blue), and 4150 km (green). Mid-ocean ridges and oceanic transforms are defined by narrow bands of shallow hypocenter earthquakes. Spreading centers and approximate full spreading rates: Southern East Pacific Rise (SEPR, B140 mm yr 1), Northern East Pacific Rise (NEPR, B110 mm yr 1) Pacific–Antarctic Ridge (PAR, B65 mm yr 1), Galapagos Spreading Center (GSC, B45–60 mm yr 1), Chile Rise (ChR, B50 mm yr 1), Northern Mid-Atlantic Ridge (NMAR, 25 mm yr 1), Southern Mid-Atlantic Ridge (SMAR, B30 mm yr 1), Carlsberg Ridge (CaR, B30 mm yr 1), Central Indian Ridge (CIR, B35 mm yr 1), Southwest Indian Ridge (SWIR, B15 mm yr 1), Southeast Indian Ridge (SEIR, B70 mm yr 1), Kolbeinsey/Mohns Ridges (KR, B15–20 mm yr 1), Reykjanes Ridge (RR, B20 mm yr 1), Juan de Fuca and Gorda Ridges (JdFR/GR, B60 mm yr 1). Earthquake data from http://earthquake.usgs.gov.
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fundamental in efforts to understand the seafloor spreading system within an integrated framework.
Methods of Monitoring Seismicity Seismicity in the oceanic ridge-transform environment is monitored using a combination of hydroacoustic technologies, which record water-borne acoustic phases associated with submarine earthquakes, and traditional seismic sensors, which record ground motion induced by compressional and shear waves propagating through the solid Earth (Figure 2). These technologies should be viewed as complementary, as each presents certain advantages and limitations. Seismometers
There are hundreds of seismometers deployed on the continents and islands across the globe. These instruments, operated primarily by governments and universities, form networks of sensors that can be used to monitor seismicity, as well as clandestine nuclear tests. However, for spreading centers that lie thousands of kilometers from the nearest seismic station, our ability to detect and locate earthquakes remains limited. Within the most remote ocean basins, global seismic networks consistently detect only earthquakes larger than roughly magnitude (M) 4.5–5, or ruptures having an approximate physical
scale of Z1 km. The signals generated by smaller events typically attenuate to below background noise levels as they traverse long solid-Earth paths between the source region and land-based sensors. Seismometers may also be deployed on or beneath the seafloor in containers that protect the instrumentation from the extreme pressure of the deep ocean (Figure 2). Seafloor instruments are commonly known as ocean bottom seismometers (OBSs). They typically are deployed from a surface ship, allowed to free-fall into position, and later located using acoustic ranging techniques. Upon retrieval, an acoustic switch is remotely triggered, causing the instrument package to release a set of anchor weights and rise buoyantly to the surface in the vicinity of a waiting ship. The detection capabilities of an OBS array depend on the number and distribution of stations. Deployment of a dozen or more instruments for year-long or multiyear observations are becoming increasingly common with recent improvements in technology. OBSs deployed at very local scales, arrays with apertures of only a few kilometers, may be used to monitor small cracking events (often with Mo0) in the vicinity of hydrothermal systems. Deployments of larger-aperture arrays, tens of kilometers across, are typically used to study volcanic and tectonic processes within a spreading segment and may provide a record of many earthquakes that would otherwise go undetected by land-based seismometers.
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Hydrophones
Due to the combined dependence of ocean sound speed on pressure and temperature, much of the global ocean exhibits a low-velocity region known as the sound fixing and ranging (SOFAR) channel. Seismically generated acoustic energy may become trapped in the SOFAR, where it propagates laterally via a series of upward- and downward-turning refractions having little interaction with the seafloor or sea surface. The attenuation due to geometric spreading is cylindrical (R 1) for SOFAR-guided waves, making transmission significantly more efficient relative to solid-Earth phases that undergo spherical spreading loss (R 2). Water-borne, earthquake-generated acoustic signals were first observed on near-shore seismic stations. As these signals arrived after the faster propagating primary (P) and secondary (S) solidEarth phases, they were termed tertiary (T). Today T phases are monitored most commonly by hydrophones deployed directly within the SOFAR. The first systematic effort to use hydrophone data to produce a continuous catalog of mid-ocean ridge earthquakes began in the early 1990s when NOAA/ PMEL gained access to the US Navy’s Sound Surveillance System (SOSUS), a permanent network of bottom-mounted hydrophone arrays within the Northeast Pacific. This enabled the real-time detection of T waves generated by seafloor earthquakes and reduced the detection threshold for seismicity at the Juan de Fuca and Gorda Ridges by almost 2 orders of magnitude. The geometry of the SOSUS network also provided a better azimuthal distribution of stations than could be accomplished using land-based seismometers. The station geometry, combined with the existence of a well-defined velocity model of the oceans, also yielded significant improvements in location accuracy. Early successes using SOSUS facilitated the development of moored autonomous underwater hydrophones (AUHs) that could be used to monitor global ridge segments. In this design, the hydrophone sensor and instrument package are suspended within the SOFAR channel using a seafloor tether and foam flotation (Figure 2). These instruments have been deployed successfully along mid-ocean ridge spreading centers in the Atlantic and Pacific Oceans, as well as back-arc spreading systems of the Marianas (W. Pacific) and Bransfield Strait (Antarctica). A typical deployment consists of only six to seven instruments that monitor B201 along axis and provide a consistent record of earthquakes with M Z 2.53.0. Although regional hydrophone arrays can provide improvements in detection and location capabilities,
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relative to land-based seismic stations, it is not presently possible to extract the focal mechanism directly from the T waveform. Similarly, only relative depth estimates are acquired through measurement of the T wave rise time, defined as the time between the onset of the signal and its amplitude peak.
Global to Regional Tectonic Patterns A view of global seismic patterns (Figure 1) shows ocean basin earthquakes to be narrowly focused along the spreading axis. Such events have shallow hypocenters and focal mechanisms that dominantly indicate normal faulting on moderately dipping (B451) ridge-parallel structures. Shallow-hypocenter earthquakes also cluster tightly along the oceanic transforms that accommodate the differential motion between offset ridge segments, with the sense of motion along these conservative plate boundaries being dominantly strike-slip on subvertical structures. These patterns reflect a stress regime arising from the motion of the plates and the geometry of their boundaries (Figure 3). Transform and Segment Boundary Seismicity
In the late 1960s, focal mechanism studies of oceanic transform earthquakes provided key evidence in support of plate tectonic theory. Early physiographic maps of the oceans had outlined the mid-ocean ridge system and identified places where this submarine mountain range appeared to be offset laterally. One might infer from these observations that the two ridge segments were once aligned and that their offset represents the cumulative displacement along the connecting fault system. In contrast, the then emerging theory of plate tectonics required the Fracture zone (inactive)
Plate A
Transform Abyssal hill faulting Spreading axis Plate B Figure 3 Schematic showing the style of faulting and seismicity associated with an idealized spreading center and transform plate boundary. Double black line indicates the axis of the oceanic spreading center. Thin gray lines show normal faults along the rift flanks. Arrows indicate the direction of relative motion between the plates. Focal mechanisms shown with shaded regions indicating compressive first motions.
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opposite sense of motion along these structures in order to accommodate the process of seafloor spreading along two ridge segments that maintain a near-constant offset through time. As Sykes showed in 1967, the first-motion direction of P waves radiated during transform earthquakes clearly supports the latter model. Today the orientation and slip direction of earthquakes are determined routinely for global events larger than M 5–5.5. These data, combined with detailed morphological observations, provide a more complete picture of the structure and seismicity of oceanic transforms and higher-order segment boundaries. Many transforms are segmented, stepping in an en echelon fashion, with pull-apart basins and more commonly intra-transform spreading centers defining the segmentation. As an earthquake propagates, offsets of sufficient size may arrest the rupture, limiting its spatial extent. The maximum size of a transform earthquake is observed to be a function of slip rate, with the largest transform earthquakes occurring at slow spreading rates (o40 mm yr 1) and having sizes of roughly M 7.5. At spreading rates 4100 mm yr 1, the maximum observed magnitude drops to below M 6.5. This reflects the thermal structure of oceanic lithosphere, with a shallower depth of seismogenic rupture at faster spreading rates, and potentially a greater tendency for transform segmentation. Earthquakes with an oblique sense of slip and orientation are sometimes observed at inside corners, the region where the active transform segment meets the ridge axis (Figure 3). Such events may represent slip on curvilinear abyssal hill faults that rotate up to 15–201 as they approach the transform. In other cases, inside corner earthquakes may be better depicted as a compound rupture, where strike-slip motion along the transform and dip-slip motion along an orthogonal abyssal hill fault occur simultaneously. In these cases, the radiation pattern cannot be well described by the double-couple force model used to represent the vast majority of global earthquakes. Transform seismicity is restricted principally to the region between the two active spreading centers, as no relative motion is required across the fault beyond this region. Fracture-zone scars, however, may be traced across ocean basins and continue to represent zones of weakness. In some areas, intra-plate stresses are sufficient to reactivate these structures, infrequently generating large earthquakes. Oceanic transforms exhibit similar kinematics to continental transforms, such as the San Andreas Fault in the United States or the Anatolian Fault in
Turkey. Oceanic transforms, however, show one major difference – an inventory of earthquakes along oceanic transforms documents a dramatic ‘slip deficit’, with the majority of all transform motion occurring aseismically. A prediction of the moment release rate along a transform can be obtained as P Mo/t ¼ nmlw, where n is the relative plate velocity, m is the shear modulus of the rock, l is the length of the transform, w is the down-dip width of the rupture, and t is the time period considered. Mo is the static seismic moment, which for an individual earthquake is the product of fault rupture area (lw), mean slip, and m. The width w is generally taken to correspond to the B600–6501 isotherm, as constrained by slip inversion results and laboratorybased deformation studies (Figure 4). When the sum of the observed seismic moment on a transform is compared with the expected moment, their ratio (or seismic coupling coefficient, a) is typically B1 for continental transforms, but much less than 1 along oceanic transforms. The average a for all oceanic transforms is roughly 0.25, indicating that three-fourths of all transform motion occur aseismically. Studies investigating the velocity dependence of a yield conflicting results. Such assessments are complicated by the presence of significant inter-transform variability in a even at a given spreading rate, uncertainty in the depth of seismic faulting, and the somewhat limited timescale of the observations. The most recent and detailed studies, however, suggest little-to-no systematic relationship with slip rate, or perhaps slightly lower a at fastspreading rates. Aseismic fault motions on the oceanic transforms are likely taken up by slow, creeping events. These slow ruptures are inefficient at producing seismic waves and so are termed quiet or silent earthquakes. It recently has been suggested that silent earthquakes may trigger large seismic quakes in neighboring portions of the transform that are more prone to stick-slip behavior. In this view, some seismic events on the transforms may be viewed as aftershocks of these silent earthquakes. At higher-order segment boundaries, propagating rifts and overlapping spreading centers accommodate ridge segmentation through a mechanism known as bookshelf faulting (Figure 5). Here the seafloor between the overlapping rift tips is deformed by simple shear, resulting in the rotation of the initially ridge-parallel seafloor fabric. The style of deformation is reminiscent of a stack of books on a shelf tipping over, with slippage between the ‘books’ occurring along the preexisting abyssal fault systems. Focal mechanisms indicate a strike-slip sense of motion, with one set of nodal planes trending
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700°C Figure 4 (a) Map of the equatorial Mid-Atlantic Ridge showing the Romanche and Chain transform faults. The focal mechanisms obtained in the broad-band body-wave modeling are joined to their NEIC locations (red circles). (b) The centroid depths of the earthquakes are plotted as circles with symbol size proportional to magnitude. Solid symbols represent earthquakes with the bestresolved parameters, and open symbols those with fewer data or poorer fits. All the depths are accurate to within 3 km. Isotherms are calculated by using a half-space cooling model and averaging both sides of the transform. Slip inversions of the 1994 and 1995 Romanche earthquakes suggest that the former ruptured from near the surface to 20-km depth, and the latter from 10- to 25-km depth. This is consistent with the B6001 isotherm controlling the maximum depth of seismic slip on these transforms. The 1992 Chain earthquake is the only event with a centroid within the crust, but it was large enough to have ruptured into the upper mantle. Reprinted by permission from Macmillan Publishers Ltd: Nature (Abercrombie R and Ekstro¨m G (2001) Earthquake slip on oceanic transform faults. Nature 410: 74–77), copyright (2001).
parallel to an abyssal fabric that becomes progressively rotated within the deformation zone. This mode of deformation is most common at fast to intermediate spreading rates, with well-studied examples along the southern East Pacific Rise (4120 mm yr 1), Galapagos (B60 mm yr 1), and Lau Basin spreading centers (B100 mm yr 1). A spectacular slow-spreading example, however, can be observed within Iceland’s southern seismic zone. Spreading-center Earthquakes
When the oceanic ridge systems are monitored using global seismic stations (MZB4.5) or regional hydrophone arrays (MZB2.53) a sharp contrast in the frequency of mid-ocean ridge earthquakes is observed as function of spreading rate (Figures 1 and 6). Along fast-spreading (480 mm yr 1) axial highs, the ridge displays a dearth of small- to moderatemagnitude earthquakes. In contrast, such events are
abundant along the rift valleys of a slow-spreading (o40 mm yr 1) ridge system. These first-order patterns largely reflect the variable thickness of the brittle layer, which controls the predicted moment release, and seismic coupling coefficient of the rift-bounding normal fault systems. The predicted rate of seismic moment release can be estimated in a manner analogous to that described P for oceanic transforms: Mo/t ¼ nmlw/sin ycos y, where y is the fault dip, n is the full spreading rate, l is the length of the plate boundary, and w is the thickness of the seismogenic lithosphere. A comparison with the observed moment release at a slowspreading rift zone indicates that only 10–20% of the plate divergence is accommodated by seismic processes. This estimate agrees well with the amount of brittle strain accommodated by the normal fault populations of the abyssal plains. Hence, riftbounding normal fault systems at slow-spreading centers display a high-level of seismic coupling, with
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180° 30′
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Figure 5 Central Lau Basin propagator with focal mechanics solutions. The Central Lau Basin Spreading Center is propagating to the south as spreading ceases along the northern tip of the Eastern Lau Spreading Center. Inset shows idealized case of bookshelf faulting, with progressive clockwise rotation of the abyssal fabric as the left-offset-propagating rift advances to the south. Focal mechanisms adapted from Wetzel et al. (1993) and the Global Centroid Moment Tensor Project (http://www.globalcmt.org). From Wetzel LR, Wiens DA, and Kleinrock MC (1993) Evidence from earthquakes for bookshelf faulting at large non-transform ridge offsets. Nature 362: 235–237.
the ratio of seismic fault slip to expected fault slip being close to 1. However, the emplacement of dikes along or near the rift axis takes up the majority (80–90%) of the plate separation. At fast-spreading ridges, commonly less than 1% of the predicted moment release is observed seismically along the rift. While most of the plate separation is accommodated by diking, as it was at slower rates, faulting studies indicated the fastspreading lithosphere undergoes an extension of B4–8%. The observed moment release is insufficient to account for this. Fast-spreading normal fault systems, therefore, appear to have low a and must accrue displacement during aseismic slip events or by
abundant microseismic activity too small to be detected by the hydrophones or global seismic stations. At intermediate spreading rates, the density of seismic events shows a first-order correspondence with the morphology of the ridge, with small- to moderate-magnitude earthquakes being abundant along rifted-spreading centers and comparatively rare at intermediate-rate axial highs (Figure 7). Intermediate-spreading-rate ridges can therefore assume both the morphologic and seismic characteristics of the fast- and slow-spreading end members. The maximum hypocentral depth of a rift zone earthquake decreases with increasing spreading rate, reflecting the temperature limits of brittle faulting
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and steeper thermal gradients at faster spreading rates (Figure 8(a)). OBS studies define additional intra- and inter-segment patterns. Within some slowspreading segments, the along-axis depth of microseismicity shallows near the center of a segment and deepens near its ends. This reflects the presence of a hotter lithosphere near segment center and is consistent with a three-dimensional (3-D) pattern of upwelling and melt focusing at slow-spreading rates. Studies on the Mid-Atlantic Ridge also show a positive correlation between the relief of the median valley and the maximum depth of microseismicity detected using OBSs (Figure 8(b)). This is consistent with stretching models that indicate greater fault offsets in regions of thicker brittle lithosphere.
Magmatism and Diking 1999−2003
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Figure 6 Mid-ocean ridge seismicity viewed from regional hydrophone arrays. (a) Mid-Atlantic Ridge (B25 mm yr 1), (b) East Pacific Rise (B110 mm yr 1) and Galapagos Spreading Center (B45–60 mm yr 1). Events located using four or more hydrophones are shown. M Z 2.5 3 earthquakes are consistently detected by these arrays. Yellow stars indicate position of hydrophones in each array. Red arrows mark firstorder (transform) offsets of the ridge axis. Data from http:// www.pmel.noaa.gov.
Prior to an eruption, mid-ocean ridge melts accumulate in crustal-level chambers. At fast-spreading rates axial magma lenses are ubiquitous, being found beneath 460% of the axis that has been surveyed using multichannel reflection techniques. At slowspreading rates, crustal-level melt is absent beneath much of the ridge axis and melt bodies may be ephemeral or localized beneath axial volcanoes. At intermediate rates, the distribution of melt can range between the two extremes, with some axial highs showing nearly continuous along-axis magma chambers reminiscent of fast-spreading centers. The accumulation of melt within the crustal chamber elevates its pressure relative to the surrounding host rock. This inflation deforms the surrounding crust and triggers earthquakes within the vicinity of the chamber. For magma to leave the chamber, the pressure (P) within must exceed the sum of the minimum confining stresses (s3) at the chamber boundary and the tensile strength (T) of the rock. The dike’s orientation will be orthogonal to the least compression stress direction and therefore it should be subvertical and aligned parallel to the axis of spreading. Although cracking in the vicinity of the dike tip creates many small earthquakes, most events of sufficient size to be detected on global seismic or regional hydrophone arrays are thought to occur on preexisting fault surfaces. In the region above the dike, a narrow zone of ridge-normal extension exists where seismicity is localized (Figure 9). A broader zone of extension exists beyond the along-axis tips of the intrusion, where seismicity will be triggered in front of a laterally propagating dike. As the dike passes, the lithosphere will be compressed at depth within the region adjacent to the dike.
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Figure 7 (a) SOSUS-detected earthquakes along the Gorda and Juan de Fuca Ridges in the Northeast Pacific, 1993–2005. Earthquakes located using four or more hydrophones are shown. M Z 2.5 earthquakes are consistently detected by SOSUS. (b) Enlarged view of portions of the Gorda and Juan de Fuca Ridge crests. (c) Cross-axis profiles from these regions.
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Figure 8 (a) Centroid depth of mid-ocean ridge earthquakes obtained from body waveform inversion of teleseismic arrivals vs. halfspreading rate. Maximum depth of seismic faulting is inferred to be less than twice the maximum centroid depth. (b) Relationship between cross-axis relief and maximum depth of seismicity. Maximum depth of seismicity was inferred from the focal depths of inner valley floor earthquakes, as recorded by OBS studies. (a) Reprinted by permission from Macmillan Publishers Ltd: Nature (Solomon SC, Huang PY, and Meinke L (1988) The seismic moment budget of slowly spreading ridges. Nature 334: 58–61), copyright (1988). (b) Reproduced from Barclay AH, Toomey DR, and Solomon SC (2001) Microearthquake characteristics and crustal VP/VS structure at the Mid-Atlantic Ridge, 351 N. Journal of Geophysical Research 106: 2017–2034 (doi:10.1029/2000JB900371), with permission from American Geophysical Union.
MID-OCEAN RIDGE SEISMICITY
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by dike Figure 9 Cartoon showing regions of compression (white) and extension (shaded) associated with a vertically propagating dike. Thin black lines indicate normal faults, with ticks on the down thrown block. Seismicity (black dots) is promoted within the extensional region and inhibited elsewhere. The width of the failure zone scales with the depth to the top of the dike.
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Much of what we know about mid-ocean ridge dike injection comes from SOSUS observations of intermediate-rate-spreading centers in the NE Pacific. On 26 June 1993, soon after monitoring began, an intense episode of volcanic seismicity was recorded in real time along the CoAxial Segment of the Juan de Fuca Ridge (JdFR), marking the first observation of such an event (Figure 10). The seismic swarm started near 461 150 N and migrated northward during the next 40 h to B461 360 N, where majority of earthquake activity occurred during the following 3 weeks. Earthquakes propagated NNE at a velocity of 0.370.1 m s 1. The reservoir that acted as the dike’s source likely resided beneath, or to the south of, the initial swarm of earthquakes. The character and propagation velocity of this earthquake swarm were very similar to dike injections observed at Krafla and Kilauea Volcanoes. Seismicity went undetected by land-based seismic networks, suggesting earthquakes Mr4.0. Since the CoAxial activity, six additional magmatic episodes have been observed hydroacoustically on the Northeast Pacific spreading centers, and several distinctive characteristics have been identified: (1)
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Figure 10 (a) Bathymetric map of the CoAxial segment of the JdFR. Circles show the locations of swarm recorded during a 23-day period of intense activity. (b) Along-segment position of epicenters vs. time during the swarm. (c) T wave rise time vs. time during the swarm. Events show a decrease in rise time consistent with the shallowing of hypocenters as the dike approaches the seafloor, erupting fresh lava in the north. Reproduced from Dziak RP, Fox CG, and Schreiner AE (1995) The June–July 1993 seismo-acoustic event at CoAxial Segment, Juan de Fuca Ridge: Evidence for a lateral dike injection. Geophysical Research Letters 22: 135–138, with permission from American Geophysical Union.
MID-OCEAN RIDGE SEISMICITY
the sequences are swarms, lacking a dominant event, with a temporal history that cannot be described by a power-law decay in event rate (i.e., Omori’s law); (2) the total number of earthquakes during swarms exceeds 10% (4350 events/week) of the variance of long-term background JdFR–Gorda Ridge seismicity; (3) swarms have several episodes of intense activity that reach 50–100 earthquakes/hour; (4) swarms last from several (45) days to several weeks; (5) earthquakes may migrate up to tens of kilometers alongaxis following the lateral injection of magma through the crust; and (6) swarms may be accompanied by continuous, broad-band energy (3–30 Hz) interpreted as ‘intrusion tremor’, resulting from magma breaking through the crust. Autonomous hydrophone recordings from the north-central Mid-Atlantic Ridge (B25 mm yr 1) indicate that diking events are less frequent, consistent with the lower rates of plate separation. During more than 5 years of monitoring, only one probable volcanogenic swarm was detected along B2500 km of ridge axis. This activity was within the Lucky Strike segment (371 N), somewhat outside of the AUH array. Earthquake locations could be determined for 147 hydrophone-detected events, with 33 of sufficient size to be located by land-based seismometers (M 3.6–5.0). In terms of total moment release, this earthquake sequence was one of largest on the Mid-Atlantic Ridge in the last three decades. The activity displayed a swarm-like temporal
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behavior and was accompanied by intrusion tremor; however, it was of short duration (o2 days) relative to previous episodes detected on the Northeast Pacific spreading centers. By far the largest episode of volcanogenic seismicity recorded on a mid-ocean ridge system occurred within the Arctic Ocean basin on the ultra-slowspreading (o5 mm yr 1) Gakkel Ridge in 1999. Seismic activity began in mid-January and continued vigorously for 3 months, with a reduced rate of activity continuing for 4 additional months (Figure 11). In total, 252 events were large enough (M44) to be recorded on global seismic networks (no hydroacoustic monitoring was in place). Subsequent geophysical studies of the Gakkel indicate the presence of a large, recently erupted flow and a volcanic peak directly in the area of seismic activity, with several large central volcanoes spectacled throughout an otherwise magma starved and heavily sedimented rift system. The duration of the 1999 activity suggests intrusion event(s) spanning several months and indicates a significant volume of magma beneath these central volcanoes. Because of their small magnitudes, earthquakes generated by diking events along fast-spreading ridge crests may be difficult to detect using hydroacoustic or global seismic techniques. Despite the greater rate of plate separation, during an 8-year period of hydroacoustic monitoring on the East Pacific Rise (101 S–101 N), only a handful of short-duration
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Figure 11 Seismovolcanic activity on the ultra-slow-spreading Gakkel Ridge near 851 N, 851 E. (a) Three-dimensional view from west (bottom) to east (top). The dark, reflective terrain is centered about a close-contoured high having a maximum vertical relief of 500 m. Red circles show the locations of epicenters, Jan.–Sep. 1999. (b) Histogram showing progression of the swarm through time, with each bar representing the number of events per day. Inset figure shows the cumulative number of events through time. There is a clear decrease in the rate of activity on 15 April (dot-dash line). Dashed line shows 6 May, when the USS Hawkbill passed over the area collecting the side scan imagery shown in (a). (a) Reprinted by permission from Macmillan Publishers Ltd: Nature (Edwards MH, Kurras GJ, Tolstoy M, Bohnenstieh DR, Coakley BJ, and Cochran JR (2001) Evidence of recent volcanic activity on the ultra-slow spreading Gakkel Ridge. Nature 409: 808–812), copyright (2001).
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(hours to days) swarms containing a few tens to hundreds of locatable earthquakes were detected. They were primarily located along intra-transform spreading centers and near segment ends, where a somewhat colder, thicker lithosphere might be expected to support larger ruptures. In January of 2006, a submarine diking event and seafloor eruption were captured for the first time by a network of OBSs. The eruption was located within the well-studied 91500 N region of the East Pacific Rise. As part of a multidisciplinary study, a network of up to 12 OBSs was deployed at this site in October 2003. Analysis has shown a gradual ramp-up in activity during a more than 2-year interval prior to the eruption, with activity culminating in an intense period of seismicity that lasted B6 h on 22 January 2006. Although many of the OBS instruments were destroyed by the eruption, the event was also recorded by regional AUHs, and the joint analysis of these data is expected to further illuminate the process of dike injection at a fast-spreading ridge. It should be emphasized that these somewhat limited observations may not be characteristic of all eruptive activity at mid-ocean ridges. Although ridge scientists often refer to a volcano-tectonic cycle, given a sufficiently long period of observation it is likely that mid-ocean ridge eruptions, like their continental equivalents, will display a power-law scaling with respect to size and duration, with smaller episodes occurring more frequently than larger ones. The scaling exponent and maximum eruption size, however, should be expected to vary as a function of spreading rate, reflecting variability in the dynamics of each system and the size of its crustal reservoir. Likewise, eruption frequency is likely to show a long-term clustering, rather than the periodic behavior sometimes alluded to.
field, strain rates from cooling are estimated to be on the order of 10 6 yr 1, about 3 orders of magnitude higher than tectonic strain rates. In 1995, a vent-scale OBS study lasting 105 days was conducted at 91 500 N on the East Pacific Rise, where a series of temperature probes also were deployed within the high-temperature vent systems (Figure 12). This study recorded a swarm of 162 microearthquakes during a period of about 3 h. Hypocenters defined a subvertical column at a depth of 0.7–1.1 km (Figure 12). Four days later, temperature sensors within a high temperature (Bio9) vent began to record increasing fluid temperature of B1 1C/day, with this trend continuing for 7 days. Vent temperature then returned to background levels during the next B120 days. Comparisons with vent temperatures during more than 3 years of monitoring showed this postseismic fluctuation to be the largest recorded, making a strong argument that the earthquakes had perturbed the hydrothermal system. It was suggested that fluids rapidly penetrated the new fractures and extracted heat from the fresh rock. A decade later, multiyear OBS monitoring, in conjunction with in situ temperature and chemical and biological studies, has been reestablished in the 91 500 N region. The results of this ongoing work, combined with similar vent-scale OBS studies conducted at intermediate- and slow-spreading vent sites, show microseismicity to be ubiquitous in hydrothermal areas. Moreover, they suggest a more complex relationship between seismic activity and vent hydrology than could have been envisioned based on the 1995 study, with temperature and flow responses exhibiting variable amplitude, sign, and phase, and many swarms producing no detectable perturbation at the vent sites.
Hydrothermal Seismicity
Tidal Triggering
Within the mid-ocean ridge hydrothermal regime, earthquake activity may be triggered in response to the removal of heat, which leads to thermal contraction and subsequent cracking, or via hydrofracture when trapped pockets of fluid are heated. Such events are typically small, with rupture diameters of 1 m to tens of meters, and can only be recorded using local OBS arrays deployed within the high-temperature vent fields. Within the shallow crust the predicted mode of failure is for purely tensional (mode I) or mixed-mode fracture, rather than the shear failure observed for the largermagnitude spreading-center earthquakes. Given estimates of the hydrothermal heat flux within a vent
Solar and lunar tidal forces exert short period stress variations that act on the Earth. Beneath the oceans and in coastal areas, stress changes influencing earthquake occurrence reflect the combined effect of the direct Earth tide and indirect loading of the Earth by the ocean tides. Together these changes are on the order of 103 104 Pa, much less than the average stress drop of an earthquake or the strength of crustal rocks. Many studies have examined the role of tides in triggering seismicity around the globe. With the exception of some terrestrial volcanic areas, earthquakes and tides are generally not correlated or weakly correlated during periods when tidal fluctuations are at their largest amplitude.
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Figure 12 (a) Ridge normal cross section of relocated microearthquakes associated with a 1995 swarm beneath the 91 500 N hydrothermal site on the East Pacific Rise. (b) Correlation between the microearthquake swarm and the fluid exit temperature at vent Bio9. Reproduced from Sohn RA, Hildebrand JA, and Webb SC (1999) A microearthquake survey of the high-temperature vent fields on the volcanically active East Pacific Rise (91 500 N). Journal of Geophysical Research 104: 25367–25378 (doi:10.1029/ 1999JB900263), with permission from American Geophysical Union.
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Julian days, 1995 Figure 13 (a) Fifteen-day time histogram of the earthquake count in 2-h bins with the ocean tide time series superimposed. Earthquakes are triggered at periods of low ocean tide, when the seafloor is unloading and stress within the crust is most extensional. Data collected during an ocean bottom seismic experiment near 481 N, 2311 E on the Endeavour Segment of the intermediate-rate JdFR. Reproduced from Wilcock W (2001) Tidal triggering of microearthquakes on the Juan de Fuca Ridge. Geophysical Research Letters 28: 3999–4002 (doi:10.1029/2001GL013370), with permission from American Geophysical Union.
Similarly, analysis of hydrophone-derived and global seismic catalogs does not show a robust correlation between tidal phase and the occurrence times of small- to moderate-size mid-ocean ridge earthquakes. However, microseismicity (Mo2) within axial hydrothermal systems commonly shows a strong correlation, with earthquakes occurring
during periods of peak extensional stress. In some locations, such as the JdFR in the Northeast Pacific Ocean, the amplitude of the ocean tides is large and these peak extensional periods correspond to times of low tide, when the seafloor is unloaded as the height of the overlying water column reaches a minimum (Figure 13). In other areas, such as the
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91 500 N region of the East Pacific Rise, tidal changes in the height of the sea surface are minimal and the Earth tide stresses have been shown to dominate. One explanation for these observations comes from laboratory studies of rock failure, which indicate that earthquake triggering in response to a periodic loading is dependent on the period of the oscillation relative to the time it takes a slip event to nucleate. Earthquake nucleation time, which is inversely proportional to stressing rate, is estimated to be on the order of a year or more for typical tectonic settings. Consequently, these systems are largely insensitive to tidal stress changes of semi-diurnal frequency. In volcano-hydrothermal systems, however, cooling- and magma-induced stress changes elevate the stressing rate by several orders of magnitude. Therefore, earthquake nucleation times become
comparable to tidal periods, and microseismicity in this setting becomes susceptible to tidal influence.
Impacts on Hydrothermal Systems Earthquakes may disturb the hydrology of marine hydrothermal systems through several mechanisms. As described previously, microearthquakes occurring beneath a vent field can alter the local permeability, creating new fluid pathways that allow the migration of cold water into the reaction zone or redirect the escape of buoyant, heated waters. Since the precipitation of minerals within the up- and downflow zones is predicted to reduce permeability over time, earthquakes events may be critical in maintaining the longevity of hydrothermal systems.
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Figure 14 Map showing acoustically derived locations of foreshocks (red dots), mainshock (yellow star), and aftershocks (white dots) of 1–7 Jun. 2000 earthquake sequence on the Western Blanco Transform. Error bars on SOSUS earthquake locations represent one standard error. Note difference in location relative to NEIC seismic epicenters (black squares), which are scattered across the plate interior. Locations of Vent 1 and Plume hydrothermal vent sites along southern JdFR are also shown. (b) Twenty-three-month temperature record at Plume hydrothermal site. Fluid at Plume vent decreased 4.9 1C during the 18 h immediately following mainshock. Three additional temperature drops appear to be associated with much smaller (MB4.0 4.2) earthquakes along the western Blanco (blue dots in (a)). (b) Eleven-month temperature record at Vent 1 hydrothermal site shows a possible delayed response to the June 2000, M 6.2 earthquake. After 7 months of steady readings (27273 1C), probe temperature decreased slightly prior to the event, but then dropped markedly by 18 1C begining 7 days after the event. Reproduced from Dziak RP, Chadwick WW, Jr., Fox CG, and Embley RW (2003) Hydrothermal temperature changes at the southern Juan de Fuca Ridge associated with MW 6.2 Blanco Transform earthquake. Geology 31: 119–122, with permission from the Geological Society of America.
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Earthquakes also induce static stress changes, with amplitudes that decay with distance from the earthquake source and become insignificant at ranges beyond 1–2 rupture lengths. These changes trigger secondary earthquakes (aftershocks) and dilate (or compress) the surrounding lithosphere. At more remote distances, up to several tens of rupture length from the epicenter, the transient stresses associated with passing seismic waves also may impact the hydrothermal system. In these cases, earthquakeinduced ground shaking is thought to jar open some hydrothermal conduits and seal others. Observations in the marine environment indicate that seismicity may create fluctuations in vent flow, temperature, and chemistry with variable characteristics. For example, drops in temperature at vent sites on the southern JdFR have been observed to correlate with earthquakes on the Western Blanco Transform at distances 430 km (Figure 14). The sensitivity, phase, and amplitude of these changes, however, vary significantly even between closely spaced sites (Figures 14(a) and 14(b)). Given the distances involved, temperature changes correlated with moderate-size (MB4.0) earthquakes on the Blanco Transform likely reflect the systems response to dynamic stress transients, while much larger earthquakes (M46) also induced significant static changes. Importantly, earthquake-induced changes within mid-ocean ridge hydrothermal systems can dramatically affect the biological communities that derive their energy from chemosynthetic processes. For example, the sub-seafloor microbial community is extremely sensitive to temperature, oxygen content, pH, and other environmental variables. It thrives in subsurface zones where the hot hydrothermal fluid mixes with entrained seawater. Changes in crustal fluid temperatures of only a few degrees, or a minor alteration of crustal permeability, can cause the prevailing microbial species to weaken and encourage new species to thrive and become dominant. Similarly, changes in the effluent thermal flux of hydrothermal vents can enhance heat output changing the size and temperature of the thermal boundary layer, a zone just above the axial floor that supports abundant and diverse macrofaunal communities. Still more dramatic impacts may be associated with the emplacement of magma at depth or the eruption of lava onto the seafloor. The heat supplied by these systems feeds massive bacteria blooms (floc) and may generate event megaplumes within the water column, huge volumes of hydrothermal fluid enriched in reduced chemicals that rise up to 1 km above the seafloor.
See also Acoustics, Deep Ocean. Hydrothermal Vent Fluids, Chemistry of. Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Seismic Structure. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Seamounts and Off-Ridge Volcanism.
Further Reading Abercrombie R and Ekstro¨m G (2001) Earthquake slip on oceanic transform faults. Nature 410: 74--77. Barclay AH, Toomey DR, and Solomon SC (2001) Microearthquake characteristics and crustal VP/VS structure at the Mid-Atlantic Ridge, 351 N. Journal of Geophysical Research 106: 2017--2034 (doi:10.1029/ 2000JB900371). Bird P, Kagan YY, and Jackson DD (2002) Plate tectonics and earthquake potential of spreading ridges and oceanic transform faults. In: Stein S and Freymueller JT (eds.) Geodynamics Series 30: Plate Boundary Zones, pp. 203--218. Washington, DC: American Geophysical Union. Boettcher MS and Jordan TH (2004) Earthquake scaling relations for mid-ocean ridge transform faults. Journal of Geophysical Research 109: B12302 (doi:10.1029/ 2004JB003110). Dziak RP, Chadwick WW, Jr., Fox CG, and Embley RW (2003) Hydrothermal temperature changes at the southern Juan de Fuca Ridge associated with MW 6.2 Blanco Transform earthquake. Geology 31: 119--122. Dziak RP, Fox CG, and Schreiner AE (1995) The June–July 1993 seismo-acoustic event at CoAxial Segment, Juan de Fuca Ridge: Evidence for a lateral dike injection. Geophysical Research Letters 22: 135--138. Edwards MH, Kurras GJ, Tolstoy M, Bohnenstieh DR, Coakley BJ, and Cochran JR (2001) Evidence of recent volcanic activity on the ultra-slow spreading Gakkel Ridge. Nature 409: 808--812. Fox CG, Matsumoto H, and Lau T-KA (2001) Monitoring Pacific Ocean seismicity from an autonomous hydrophone array. Journal of Geophysical Research 106: 4183--4206. Sohn RA, Hildebrand JA, and Webb SC (1999) A microearthquake survey of the high-temperature vent fields on the volcanically active East Pacific Rise (91 500 N). Journal of Geophysical Research 104: 25367--25378 (doi:10.1029/1999JB900263). Solomon SC, Huang PY, and Meinke L (1988) The seismic moment budget of slowly spreading ridges. Nature 334: 58--61. Sykes LR (1967) Mechanism of earthquakes and nature of faulting on the mid-oceanic ridges. Journal of Geophysical Research 72: 2131--2153. Tolstoy M, Cowen JP, Baker ET, et al. (2006) A seafloor spreading event captured by seismometers. Science 314: 1920--1922 (doi:10.1126/science.1137082).
MID-OCEAN RIDGE SEISMICITY
Wetzel LR, Wiens DA, and Kleinrock MC (1993) Evidence from earthquakes for bookshelf faulting at large nontransform ridge offsets. Nature 362: 235--237. Wilcock W (2001) Tidal triggering of microearthquakes on the Juan de Fuca Ridge. Geophysical Research Letters 28: 3999--4002 (doi:10.1029/2001GL013370).
Relevant Websites http://www.globalcmt.org – Global CMT Web Page.
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http://www.pmel.noaa.gov – National Oceanic and Atmospheric Administration’s Acoustic Monitoring Program. http://earthquake.usgs.gov – US Geological Survey Earthquake Hazards Program.