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FLED Project Proposal

I. Summary

We propose to install 25 broadband seismometers between Florida and Edmonton to create (with existing permanent stations) a 35-station quasi-linear array (Figure 2). Stations, borrowed from the IRIS PASSCAL program, would operate for 15 months starting Spring, 2001. Data (3-channel, 40 samples per second) would be processed into SEED volumes at Washington University using the PASSCAL Database, and made publicly available through the IRIS DMC. Data analysis, at Brown and Washington Universities, would be primarily aimed at studying several structural aspects of (1) the core-mantle boundary (CMB) and D'' beneath the central and circum-Pacific, and (2) the upper mantle and crust beneath North America.

II. Why a linear array?

Linear arrays can maximize the scientific return from a finite number of instruments, as a higher resolution is attained than if the stations are deployed in a 2-D pattern. Combining 25 portable PASSCAL broadband sensors with 10 permanent IRIS-GSN, CNSN, and SLU-USNSN stations, we propose to have 35 seismometers distributed over 3750 km (33.8°), providing us with a powerful image of a long 2-D slice into the Earth. Many CMB features could be resolved with body waves at a resolution previously not attained, with coverage still large enough to identify variations in spatial trends. Surface-wave modeling would span a large enough distance to provide a continent-sized reference frame for studies of the deep structure of continents, yet have enough stations to resolve variations across different N. America terranes. Existing N. America stations would provide information on off-strike structure and regional seismicity.

Stations need not be exactly along-strike (allowing use of the 10 permanent broadband stations). Inter-station spacing need also not be even, allowing an array design that is the best compromise between length and resolution. Station density is greatest crossing the Midcontinent Rift in Iowa, but is also denser crossing the Reelfoot Rift and the southeastern terrane boundaries (Figure 3). Station densities were chosen while maintaining the required spacing between other stations for the surface wave and CMB studies.

III. Why this particular location and orientation?

Passive seismic experiments are gambles, waiting for earthquakes to occur, but the odds are improved by designing deployments to take maximum advantage of the earthquakes that are the sources for the experiment. FLED is along the trend of the seismogenic zones of Alaska, the Aleutians, Kamchatka, the Kurils, Japan, Taiwan, the Philippines, and parts of Indonesia (Figure 4). Passive seismic experiments should come with the same fine print that appears in Mutual Fund advertisements ("Past performance is not necessarily an indication of future yields"), but our odds are very good of having many usable seismic sources. The line of seismicity from Alaska to Indonesia also exists at many distances that would provide a variety of seismic phases needed for investigating the top and bottom of the mantle. Table 1 lists the epicentral distance ranges required for the seismic phases discussed in the following science sections. We also list the along-strike seismogenic zones and the distance ranges at which example earthquakes would be recorded across FLED. These are not the only distance ranges available: as Figure 4 shows, the Alaska-Indonesia seismicity is nearly continuous.

There are other highly seismically-active regions (Table 1 - Tonga, Kermadec, Vanuatu, New Guinea) that will provide excellent sources for some of the CMB studies. If we select a site mid-way along the proposed FLED array (St. Louis) as an example, based on previous history it will record half (49.6%) of all seismicity in the 100° - 135° range that is required for many of the CMB phases like Sdiff, Pdiff, and SKS [Wysession, 1996c]. The only other continental region better for this is South Africa.

Large earthquakes along the FLED great circle path can occur in other regions: (1) New Madrid , (2) El Pilar fault, Venezuela (a source of surface waves from the south), and (3) a seismogenic segment of the Southeast Indian Ridge (reaching FLED at 158° - 192°). The antipode of this SEIR segment is where FLED crosses the Midcontinent Rift with a higher density of stations. Events from these three locations are long shots (we are not holding our breath), but they would be a special bonus if they occurred.

The other justification for the location and orientation of the proposed array is its location relative to the lithospheric components of North America (Figure 3). FLED would cross 3 different Archean cratonic fragments, two major rift zones, and a host of orogenic and accretionary terranes. The trends of many of these tectonic lineations and terrane boundaries are at a high angle to the trend of the FLED array. Body and surface wave studies will provide a slice of the deep lithospheric structure from the oldest Archean fragments to the most recent allochthonous additions. We have an idea of the deep continental structure from large-scale inversions, in particular the fast (presumably cold) and deep (>250 km) cratonic root [Grand, 1994; van der Lee and Nolet, 1997a; Grand et al., 1997; Larson et al., 1998], but our understanding is largely preliminary without seismometers within the regions. For instance, the Midcontinent Rift is one of the best examples of a large failed continental rift to be found on Earth, and it happens to be in our own backyard, but we have no knowledge of what lies deep beneath the rift, as no broadband seismometers have ever been deployed across it.

IV. Relation to other deployments

North America is not bereft of broadband seismometers, though it is true that a large gap exists between Cathedral Caves (CCM) and Edmonton, and that if Flin Flon goes down, there is a 3000 km-diameter circle within North America (CCM to Yellowknife, and Corvallis to Frobisher Bay) without a single IRIS GSN or FDSN station. But as seen in Figure 2, many USNSN stations exist in the United States, and as of Spring, 1998, their data are being archived at the IRIS DMC. While data are not currently usable because of header inconsistencies, data-less SEED volumes are planned to be constructed by the Albuquerque USGS group, and these data should certainly be available by 2001. These stations are also 40 sps, and easily incorporated into our analyses. There are also several broadband Canadian stations part of the Canadian National Seismic Network (CNSN) with publicly-available data. We plan that ten of the permanent stations will be part of our linear array. Many of these are operated by R. Herrmann (St. Louis University) as part of the Cooperative New Madrid Seismic Network (SLM, FVM, SIUC, PVMO, PWAL, UTMT). We are in regular communication with Dr. Herrmann, who is committed to providing the necessary SEED-volume quality control, and to upgrading sensors from CMG-40T's to 3ESP's or 3T's. Additional off-axis stations will play an important role concerning 3D aspects of the analysis: CMB-phases, surface waves, and regional sources, and are an additional benefit of doing such a cross-continental deployment in North America.

The sparseness of mid-North American broadband stations has been the subject of discussion within the seismological community, with a "USArray" proposed to increase station density [Ekström and Sheehan, 1998]. Discussions, within the IRIS GSN Committee and elsewhere, are in their very early stages, and proposals include permanent stations and/or a moving grid of seismometers, like that proposed by Humphreys et al. [1998]. These demonstrate the interest in having better seismic coverage within the North American interior. However, none of the discussions have involved an array resembling FLED - 35 stations linearly spanning 34°. And at best, it would be 4-5 years before funding could be obtained for such an initiative, and many years after that before science would be done. We would hope that the science learned from FLED could provide useful lessons for the construction of a USArray, should it come to fruition.

The only previous broadband array in this region was the 1989 Archean-Proterozoic Transition Experiment [Silver et al., 1993], which extended ENE from northeastern Wyoming to western Ontario (FLED crosses this transect with a perpendicular trend in southern North Dakota). Unfortunately, this data set has been underutilized because this was the first time that data were written to disk using the REFTEK units, and there were "unrecoverable errors on the order of a fraction of a second about 25% of the time" [Silver et al., 1993]. Nonetheless, this array produced excellent studies of earth structure using relative arrival times and amplitudes [Silver and Kaneshima, 1993; Bokelmann and Silver, 1993; Silver and Bina, 1993].

Refraction surveys have been carried out in some of the FLED regions, providing important shallow-structure constraints for inversions of deeper structure. For example, a survey across the Iowan MCR (and borehole) provides the velocity structure down to 15 km [Chandler et al., 1989], which can be held constant in our surface-wave inversion. The Deep Probe Active Source Experiment [Henstock et al., 1998], providing structures < 150 km deep from New Mexico to Alberta, and LITHOPROBE studies in southern Canada [Clowes et al., 1996] will allow us to "pin down" our velocity model where FLED passes into Alberta.

V. Science Outline

Many seismic investigations can be done with a 34°-long cross-continental broadband array. We plan to use the data to build upon areas of research into earth structure that is already underway at Brown and Washington Universities: (1) lowermost mantle and core-mantle boundary structure, (2) transition zone and upper mantle discontinuities, (3) upper mantle anisotropy, (4) structure of the North American lithosphere. Within these areas, our work cannot possibly be exhaustive. But it is also our hope that this data set will be used in many other ways by future studies, as has been the case with the data from MOMA.


The core-mantle boundary (CMB) is one of the most unusual parts of the Earth, one of the most important for understanding Earth's dynamic evolution, and one currently receiving much attention from the geophysical community [Lay et al., 1998b]. Tremendous achievements in understanding the structure of the lowermost mantle, D'', have been made with the global distribution of seismometers, but the large spatial aliasing causes some severe limitations. PASSCAL-type deployments, with many high-quality sensors strategically placed, provide unique windows onto the CMB, revealing structures otherwise not visible.

The AGU recently published a Geodynamics Series volume on the CMB region [Gurnis et al., 1998]. During its assembly, this volume showed that a very wide variety of geological structures must exist at the CMB to create the range of seismic observations: anisotropy, scattering due to topography or small-scale heterogeneity, large vertical velocity gradients, an ultra-low velocity layer (ULVZ), and discontinuities. There was also a sense that more questions were being raised than answered. We included "white papers" designed to recap what was known about aspects of the CMB [Garnero et al., 1998; Lay et al., 1998a; Wysession et al., 1998a], but found that it was very hard to say anything definitive. For example, compiling the results from over 40 studies of the "D'' discontinuity," there was a sense that this may not actually be a discontinuity in the same way we think of upper mantle discontinuities, but rather a complex 3D set of structures that give the appearance of a discontinuity when using 1D modeling [Wysession et al., 1998a].

One of the most important scientific contributions we hope to make is an examination of the transition from the CMB beneath the mid-Pacific to that beneath the Pacific rim. The mid-Pacific D'' displays very slow VS and VP, azimuthally-varying anisotropy and a strong ULVZ. The Pacific-rim D'' has mostly fast velocities, transverse isotropy and a narrow or non-existent ULVZ. The transition between the two is poorly understood. Is it a sharp or gradual transition? Is it the same for velocities, anisotropy and the ULVZ? Is it different for VS and VP? Our plan is to use a variety of different seismic CMB phases recorded at FLED and off-axis stations to answer these questions about the structure of D'' and the CMB. All of these studies will use techniques that we have already designed or adapted to study MOMA and global-network data, so a minimum of project time will need to be spent on coding and algorithm development.

1. Small-scale lateral velocity variations.

There has been a long-standing discussion as to the scale of heterogeneities at the base of the mantle [Su and Dziewonski, 1991; Lay, 1991]. Global inversions show significant heterogeneity at the 1000+ km scale [Dziewonski et al., 1996; Wysession, 1996b], but where there is resolution for small-scale features, they are observed [Grand et al., 1997; van der Hilst et al., 1997]. And regional array studies have identified lateral variations at scales of 100's [Weber, 1994], and 10's of km [Bataille and Lund, 1996]. We plan to examine this in several different ways, including documenting the lateral variations in VP/VS ratios. P and S velocities are actually anti-correlated in some parts of D'' [Bolton and Masters, 1996], but it has been hard to determine whether this is due to chemical variations or azimuthal anisotropy. Our simultaneous examination of both core-diffracted P and S waves will provide a strong means of looking at this.

*** Sdiff and Pdiff arrivals. Core-diffracted Sdiff and Pdiff waves can spend a lot of time within D'', and provide a good tool for looking at lateral variations. While this is often in the form of slowness estimates over large distances [Wysession et al., 1992; Souriau and Poupinet, 1994; Hock et al., 1997; Valenzuela et al., 1998b], we also looked at station-to-station variations with the MOMA data [Wysession et al., 1998c]. We plan to look at lateral variations at the scale of 100 km at the base of the mantle using a double-array system with many stations (FLED) and many earthquakes (Japan to Indonesia) all along the same great circle path. Data are band-passed filtered around a period of 25 s to avoid effects of dispersion (see #2B, below), and travel-time station corrections are formed by using events (Kuril and Aleutian) from pre-diffraction distances along the same back-azimuth. This will study the Japan-Canada CMB corridor beneath the northern Pacific rim, of particular interest as a proposed site of significant amounts of accumulated paleoslab rock [Lithgow-Bertelloni and Richards, 1998]. MOMA studies also suggest that the region has anomalously low VP/VS ratios for west-east paths (Figure 5), suggesting that patterns of flow associated with the paleosubduction are causing either a mineralogical or textural preferred orientation [Wysession et al., 1998b].
*** SKS-SKKS-S/Sdiff differential travel times. Relative times between different core phases provide good measures of lateral D'' variations because uncertainties in the rupture process and near-source and receiver structure is largely removed. This can involve SKKS-SKS [Lay and Young, 1990], SKS-S [Garnero and Helmberger, 1993], PKP branches [Creager and Jordan, 1986; Song and Helmberger, 1993], and SKS-Sdiff [Dziewonski et al., 1996; Kuo and Wu, 1997]. Because the latter involves a combination of SV (SKS) and SH (Sdiff) energy, it requires an assessment of upper mantle anisotropy beneath the FLED stations. Differential SKS, SKKS and Sdiff times are obtained with alignments of the initial waveform segments, altered to account for differential attenuation (and for SKKS, a °/2 Hilbert transform), in the manner of Wysession et al. [1995]. We plan to use a combination of SKS, SKKS and Sdiff (or S), as is being carried out with MOMA data by Valenzuela et al. [1998a], to provide a study of the lateral variations (1) beneath the Japan-Canada corridor that is complementary to that of the Sdiff double array approach, and (2) beneath the northeastern Pacific, using earthquakes from the Tonga, Kermadec, Vanuatu and New Guinea trenches.
*** PKP-Pdiff differential times. Wysession [1996b] used PKP-Pdiff times to create a global map of large-scale P-velocity variations above the CMB. We plan to do this on a regional scale using FLED data. The PKP(DF) and Pdiff arrivals are both usable in the 120° - 160° range, and we would use seismic events along the Pacific subduction zones of Sumatra-Java-Sumba-Timor-Banda-Philippines-New Guinea-New Ireland to invert for the P velocities of a small 30° x 50° region of D'' beneath the Hawaiian seamount chain. We plan to examine the transition from the slow velocities of the "Equatorial Pacific Plume Group" to the fast velocities of the "China High" [Su et al., 1994].
*** SdS and ScS, for variations in the "D'' discontinuity". The term "D''" was originally used to describe a region at the base of the mantle with velocities lower than the projected mantle adiabat, and so was associated with the presence of a thermal boundary layer (TBL), "D''" is now often defined by the inferred location of a velocity increase about 250 km above the CMB. It seems incontrovertible that there is fast-velocity rock at the top of D'', but what form this "discontinuity" takes is unclear, especially as observations of the discontinuity turns on and off over very short distances [Lay et al., 1997], and that its height above the CMB ranges from 100 - 450 km [Kendall and Shearer, 1994]. One problem is that most information comes from many earthquakes recorded at a small number of stations, so events with different source mechanisms, magnitudes, depths, and noise levels are compared to each other. By examining SdS waveform changes along a continual path of many stations, we hope to better grasp the scale of D'' discontinuity variations. For example, Wysession et al. [1997] presented an example where the SdS amplitude increased and then decreased dramatically over a 200 km horizontal CMB path beneath the Caribbean (simultaneously, ScS decreased and then increased, suggesting focussing effects of 3D heterogeneities or lateral variations in a discontinuity's seismic impedance). Our array is well suited to record SdS using earthquakes from both Kuril and South American trenches, which have provided the foundation for studies of the D'' discontinuity [Mitchell and Helmberger, 1973; Lay and Helmberger, 1983; Young and Lay, 1990; Kendall and Nangini, 1996; Matzel et al., 1996; Lay et al., 1997]. We plan to use data in the 65° - 75° range to isolate the SdS phase using a deconvolution of the instrument response, determining the depth of the apparent discontinuity from the relative S, SdS, and ScS arrival times, and also measuring the SdS/S/ScS amplitude ratios. We also plan to use reflectivity-synthetic waveform forward modeling of the combined S-SdS-ScS arrivals out to distances of 105° to further examine the lateral variation in the discontinuity.

Key Questions:

  • What is the scale of D'' heterogeneity beneath the central Pacific?
  • Beneath the circum-Pacific? Is there a difference? Is there a noticeable transition?
  • Are there velocity differences between SH, SV and P velocities in differentregions?
  • Is there a sharp transition between the central Pacific and circum-Pacific?

2. Vertical velocity structure of the base of the mantle.

Frustratingly little is known about the vertical velocity structure of D''. The SdS and PdP arrivals that image the D'' discontinuity have little resolution of the bottom of D''. But core-diffracted waves are particularly sensitive to the structure of the base of the mantle [Doornbos and Mondt, 1979a,b], and the vertical velocity structure of D'' is very important for understanding the dynamics of the D'' TBL. For example, Ritsema et al. [1997] modeled Sdiff waves to determine that the velocity gradient beneath a region in the Pacific was strongly negative over at least 200-300 km above the CMB. This might mean that very high viscosities or densities occur within the CMB to suppress instabilities that might break up the TBL, or that this TBL is a chemical boundary layer decoupled from the rest of the mantle [Wysession et al., 1998a]. We plan to use several techniques to examine the lateral variations in the vertical structure of D''.

*** SPdKS, for Ultralow Velocity Zone (ULVZ) structure. An important recent discovery has been the ULVZ just above the CMB. The primary tool for examining this layer is the phase SPdKS, an SKS wave that briefly travels as a diffracted P wave along the CMB at either or both of its core entrance/exit points [Garnero et al., 1993]. SPdKS first appears as a shoulder after SKS, and anomalously late arrivals signify the presence of a narrow layer with P velocities 10% slow. ULVZ is usually observed in regions of mantle upwelling (mid-Pacific), and therefore may be thicker there than in regions of downwelling (Pacific rim), where the layer may either be very thin or not present [Revenaugh and Meyer, 1997; Garnero et al., 1998]. SPdKS is first seen beyond about 105° - 108° when SKS energy begins partitioning into SPdKS [Silver and Bina, 1993], and has been modeled out to 125°. Data in this distance range can be observed at FLED from earthquakes in the Kermadec, Vanuatu, New Guinea, Philippines and Taiwan subduction zones. Modeling of SKS and SPdKS phases recorded at the MOMA array indicates that (1) a ULVZ exists at the CMB beneath the central-Pacific where overall D'' velocities are slow, and (2) a ULVZ is not required beneath the north Pacific rim where overall D'' velocities are fast (although the existence of a very thin (< 5 km) ULVZ cannot be ruled out) [Fischer et al., 1998] (Figure 6). We expect to resolve the transition in the ULVZ thickness from the western Pacific (thick) to beneath Japan and eastern China (thin).
*** Sdiff and Pdiff slownesses as functions of frequency. A powerful tool for examining the vertical structure of the crust and upper mantle is the geometrical dispersion of surface waves. A similar technique can be used with long-profile Sdiff and Pdiff waves. Because the velocity structure of the base of the mantle varies vertically, Pdiff and Sdiff of differing wavelengths will travel at different velocities, resulting in a form of geometrical dispersion [Okal and Geller, 1979; Souriau and Poupinet, 1994]. However, the effect is subtle, and extremely good data at many stations are required to give slowness variations above noise levels. We measured the frequency variations in Pdiff and Sdiff slownesses across the MOMA array for 17 earthquakes to a degree never before possible (Figure 7), and found that significant regional variations from a PREM structure exist laterally [Wysession et al., 1998c]. This technique can easily recognize a positive velocity gradient, and no such regions were found at the base of D''. We plan to examine the lateral changes in the thickness of the TBL and strength of the D'' discontinuity using Sdiff and Pdiff dispersion data from earthquakes in Taiwan, the Philippines, New Guinea and Indonesia.
*** Sdiff and Pdiff amplitudes as functions of frequency. As diffracted waves travel along the CMB, an anomalous decrease in amplitude will result from high intrinsic attenuation (1/Q) or a positive velocity gradient (energy turns away from the layer). But anomalously large amplitudes are an indication of negative velocity gradients (energy is trapped in the layer), as attenuation cannot cause this (even Q = ° cannot increase amplitudes enough). So diffracted wave amplitudes are a great way to identify negative velocity gradients in D''. Valenzuela and Wysession [1998] forward-modeled Sdiff amplitude variations for two different CMB regions using a variety of possible velocity structures. Both regions, beneath easternmost Siberia and the eastern Pacific, were best modeled with a discontinuous increase underlain by a strong negative velocity gradient. The very wide TBL supports the findings of Ritsema et al. [1997]. The Pacific region had a shallower discontinuity and a more severe negative gradient. We plan to use this approach for Sdiff and Pdiff arrivals from events along the FLED great circle. We can also apply a correction not possible with the MOMA data: using S and P data from pre-diffraction distances such as the Kurils to calibrate the amplitude frequency dependence due to the wave paths beneath the North American upper mantle.

Key Questions:

  • Is the ULVZ continuous? Smooth or lumpy? Visible under the north Pacific rim?
  • Is the ULVZ transition from central to circum-Pacific gradual or sudden?
  • Are there any regions with a positive velocity gradient at the base of D''?
  • How thick is the TBL? Does this vary laterally? The same for P and S waves?
  • Is the structure of D'' more complex? Does it require 3-D modeling?
  • Is a chemical boundary layer required? Beneath the circum-Pacific as well as central Pacific?

3. Anisotropy.

While the base of the mantle appears to be highly anisotropic [Lay et al., 1998a], the form of this anisotropy is not fully known. Beneath regions of paleosubduction-related downwelling, it is usually the case that horizontally-propagating SH waves travel faster than SV waves, and transverse isotropy is a good model, perhaps resulting from horizontally-layered melt or seismically-slow chemical anomalies [Kendall and Silver, 1996; Matzel et al., 1996; Garnero and Lay, 1997]. However, beneath the central Pacific results have been more complicated, and the presence of significant azimuthal anisotropy has been suggested [Vinnik et al., 1989a, 1995a, 1998; Maupin, 1994; Winchester and Creager, 1997; Pulliam and Sen, 1998; Ritsema et al., 1998; Valenzuela and Wysession, 1998b]. While many causes are possible [Karato et al., 1995; Wysession, 1996a], a reasonable mechanism for the anisotropy beneath the central Pacific is the vertical orientation of lamellar structures swept up within the broad regional upwelling [Lay et al., 1998b].

Fouch et al. [1998a] used the differential times of MOMA SHdiff and SVdiff arrivals, corrected for upper mantle anisotropy beneath the stations, to examine small-scale anisotropy variations across the sub-Pacific CMB. We found that SHdiff is consistently fast relative to SVdiff in the sampled portions of both the northern circum-Pacific and the central-Pacific, but that the magnitude of the splitting is greater in the central-Pacific. Because of the geometry of the MOMA array, it was possible to differentiate splitting between Sdiff paths and better localize where the anisotropy occurs within D''. The FLED stations would provide much better sampling of CMB anisotropy beneath the western Pacific rim and Eurasia than is possible with existing permanent stations, and the array geometry would also allow for good localization of anisotropy within D''. In addition, FLED stations would record on paths at different back-azimuths through the well-studied northern Pacific rim, enabling us to determine whether truly transverse isotropy is required, or whether anisotropy is azimuthal but previously studied over only a limited range of back-azimuths. SHdiff and SVdiff arrivals for earthquakes from Kermadec to Taiwan will better identify and reveal the transition between the different styles of central and circum-Pacific D'' anisotropy.

Key Questions:

  • Is circum-Pacific D'' anisotropy consistently transversely isotropic?
  • Does central Pacific anisotropy follow any recognizable patterns?
  • Is there a sharp transition between the two that might correspond to paleoslablocations?


Seismic estimates of the sharpness and depth distribution of mantle discontinuities are important for constraining models of Earth's composition and temperature. An advantage of imaging mantle discontinuities with data from dense portable deployments, as opposed to global networks with sparser coverage, is the ability to better track the deep thermal and chemical signatures of surface tectonic features like hotspots [Dueker and Sheehan, 1997; Yang et al., 1997], mid-ocean ridges [Yang et al., 1998], and craton/orogen transitions [Chen et al., 1997; Dueker and Sheehan, 1998; Li et al., 1998a,b]. Mantle discontinuities and the Moho beneath the FLED array will be imaged by common-depth-point binning of move-out-corrected Ps receiver functions. As with MOMA array data [Li et al., 1998a,b], crustal structure will be modeled using Moho converted phases and reverberations, with deeper converted phases corrected for crust/mantle heterogeneity using these crustal models combined with existing velocity models [e.g. Grand et al., 1997; van der Lee and Nolet, 1997a; Larson et al., 1998].

A very important question concerns the depth of cold temperatures associated with the North American mantle keel. MOMA stations crossed the eastern edge of the keel, and neither the topography of the 410-km discontinuity nor the thickness of the transition zone to the north of the array (sampling more deeply into the keel) showed significant variations correlated with the keel margin (Figure 8). These results indicate that broad thermal anomalies of more than 100\260C-150\260C associated with the keel are confined to shallower depths. However, the FLED array would sample much more deeply into the keel interior. Ps conversions at widely-spaced permanent Canadian stations show little variation in transition zone thickness [Bostock, 1996], but FLED would provide a much denser sampling of discontinuity topography, enabling us to ascertain or disprove over finer spatial scales whether cold keel material, or cold mantle downwellings beneath the keel, impinge upon the 410-km discontinuity.

Many Ps studies have found discontinuities in the 200-300 km depth range, and with MOMA they changed character and deepened across the keel margin [Li et al., 1998b]. FLED data would illuminate similar changes in discontinuities across the keel margin in the southeastern U.S., and track the depth and strength of the ∼300-km phase over more than 2000 km into the keel interior. Does this boundary correspond to a weak low velocity zone at the base of the keel, or is it internal to the keel, marking the bottom of a mechanically strong anisotropic layer, as suggested by Gaherty and Jordan [1995] in Australia, and Bostock [1997] in the Canadian Slave Craton? We will also look for evidence of other internal layering within the lithospheric keel, and how these layers might reflect the process of root accretion [c.f. Bostock , 1998].

A third and more speculative question is whether transition zone discontinuity topography can constrain the location of the subducted Farallon plate that appears as a high velocity slab beneath the eastern North American lower mantle. The sub-MOMA 660-km discontinuity showed a ∼25 km depression to the south of the western stations, roughly correlating with the transition zone position of the Farallon slab imaged by Grand [1994], Grand et al. [1997] and van der Hilst et al. [1997]. However, van der Lee and Nolet [1997b] place the Farallon slab in the transition zone further west. We hope to resolve if significant residual cold temperatures are present in the Farallon slab and if the slab is located in the transition zone beneath the central U.S. by looking for topography on the 660-km discontinuity beneath the FLED array.

Key questions:

  • How deeply do thermal anomalies associated with the continental keel extend?
  • How do upper mantle phases vary across the margin of the keel and into its interior?
  • Do they reflect structure and/or rheology within the keel or at its base?
  • Can transition zone discontinuities constrain the location of the subductedFarallon plate?


1. Variations in lithospheric roots beneath different cratons and orogens.

A) Rationale. The mantle convects in a particular manner, different than any other planet, and the history of this convection is recorded in the assembly of our continents. North America has a particularly rich geological history involving the assembly of a variety of different types of terranes over the past 4 Ga. The identification of the sub-crustal identities of these terranes can provide a better sense of this history.

However, much uncertainty exists in the deep structure of continents due to inadequate sampling. For example, it is commonly thought that Archean terranes have thinner crusts but thicker lithospheres than Proterozoic orogens [Durrheim and Mooney, 1991]: the roots of cratons may extend deeper than 250 km [Grand, 1994; van der Lee and Nolet, 1997a], whereas as the roots of orogens may only be 160 km deep (as with deep seismic profiling under the central Trans-Hudson Orogen [N\351meth and Hajnal, 1998]). But not all studies are in agreement. van der Lee and Nolet [1997a] found anomalously slow velocities 100 km deep within the Archean Wyoming Province, while the Deep Probe seismic refraction experiment [Henstock et al., 1998] found the Wyoming Province to have very fast velocities to at least 150 km in depth, as well as an anomalously thick crust (∼50 km). Variations in continental lithosphere thicknesses may have important ramifications for the future of the continent. Deep Archean roots may channel hot spot flow into the shallower roots of Proterozoic orogens, determining when and where a continent (that is already under trench roll-back subduction-related extensional forces) will rift apart [Hill, 1991]. Is this what happened with the MCR? Why did it stop rifting? If subduction starts on both sides of North America, will the Yellowstone (or another) hotspot initiate rifting? Where would it occur? Is this why Greenland rifted from Labrador along a pre-existing suture? The first step is getting a better understanding of the deep continental structure.

B) Method. Surface wave inversions for sub-array lateral velocities would be carried out using a procedure initially developed by Herrmann [1987] that we have altered to include a genetic algorithm parameter search. A 2-D grid across the array is established, with a set number of either fixed or floating-depth blocks between stations. Substation Moho depths are fixed, based upon the results of the Ps studies mentioned above (Part B). Shallow velocity structures (refraction-determined crustal velocities, Pn velocities, etc.) can also be incorporated and held constant. A set of 50 randomly generated models is selected within the allowed bounds for the grid block parameters. Synthetic seismograms at all stations are created for each model using the algorithms of Herrmann [1987]. The parameters of each model are converted into binary values and concatenated into long binary strings. The misfits (costs) between the fundamental Love and Rayleigh waves and all models' synthetic counterparts are computed, and these costs govern the likelihood that models will survive the genetic algorithm process of selection, reproduction (binary string cross- over), and mutation that create the next generation of 50 models [Koper and Wysession, 1998; Koper et al., 1998]. This process is repeated until convergence occurs. This method was recently developed to determine the shallow crustal structure of a region within eastern Maine using Rg waves [Al-eqabi et al., 1998].

There are numerous 1D models of North American mantle velocity structure (e.g., Grand and Helmberger [1984]; Iyer and Hitchcock [1989]) that will provide the basis for quantifying lateral and vertical heterogeneities. Regional refraction, reflection and surface wave studies have also been done in several regions common to FLED, providing constraints on shallow structure that can aid in the accurate recovery of deeper structure. Some of these regions include (this list is not yet complete): Georgia [Dorman, 1972; Kean and Long, 1980; Nelson et al., 1985a,b; Taylor, 1989], Alabama [Long and Mathur, 1971; Long and Liow, 1986], Tennessee [Prodehl et al., 1984; COCORP Atlas, 1988], the Reelfoot Rift [McCamy and Meyer, 1966; Austin and Keller, 1982; Ginzburg et al., 1983; Mooney et al., 1983; Braile, 1989], Missouri [Stewart, 1968], Iowa [Cohen and Meyer, 1966; Chandler et al., 1989], North Dakota [McCamy and Meyer, 1966; Warren et al., 1972], Saskatchewan [Hajnal et al., 1984; Kanasewich and Chiu, 1985; Morel-a-L'Huisser et al., 1987], and Alberta [Chandra and Cumming, 1972; Henstock et al., 1988]. In addition, there are studies of Pn waves for sub-crustal P velocities, and crustal thicknesses [e.g., Braile et al., 1989; Mooney et al., 1996]. We will not be beginning from scratch in terms of starting crustal models.

C) Station Coverage. The surface wave technique just mentioned will be used to investigate the deep structure across FLED transect. The following is a list of the major crustal and tectonic divisions.

  • Southern Hearne Province: 2.8 Ga Archean craton [Frost and Burwash, 1986].
  • Wyoming Province: Pre-3.1 Ga Archean craton [Peterman and Futa, 1988], sutured to the Hearne along the 1.9 Ga Great Falls Tectonic Zone [O'Neill and Lopez, 1985], that underlies the 300-450 Ma Williston Basin, [Leighton, 1996].
  • Trans-Hudson Orogen (Dakota segment): 1.8 Ga suture between the Wyoming and Superior Cratons. [Thomas et al., 1987].
  • Minnesota Foreland (Superior Province): 3.6 Ga Early Archean fragment. [Goldich and Fischer, 1986], overthrust 2.7 Ga ago by the Superior province along the Great Lakes tectonic zone [Hoffman, 1989].
  • Midcontinent Rift Zone (within the Penokean Orogen): 1.1 Ga continental rift [Allen et al., 1995]. within rocks of the 1.85 Ga Penokean Orogen [van Schmus et al., 1996].
  • Southern Central Plains Orogen: 1.7-1.6 Ga accreted tectonic belt [van Schmus et al., 1996].
  • Eastern Granite-Rhyolite Province: 1.51-1.43 Ga granites and rhyolites generated from partial melting of older lower crust [Nelson and DePaolo, 1985]. Includes the 0.6 Reelfoot Rift zone associated with the break-up of Rodinia [Hoffman, 1989].
  • Grenville Province: 1.3-1.0 Ga imbricated accreted crust that formed during 1.3-1.0 Ga [Hoffman, 1989].
  • Appalachian Fold Belt: Late Proterozoic rift-related igneous and sedimentary rocks deformed during 400 - 350 Ma Europe-Africa orogeny, overlying Grenville basement [Rankin et al., 1988].
  • Tallahassee/Suwannee Coastal Terrane: 0.5-0.8 Ga accreted terranes, likely Pan-African and Avalonian [Williams and Hatcher, 1982; Dallmeyer, 1987; Rast, 1989].

2. Deep Structure of the Midcontinent Rift.

A focus of our lithospheric studies would be in Iowa, where FLED crosses the Midcontinent Rift (MCR), and we have increased station density in this region. The MCR is a failed rift that was tectonically active during 1.108 - 1.087 Ga [Allen et al., 1995], and is associated with a 100 mGal gravity high. Reflection and refraction studies have crossed the MCR (two published from Iowa: Cohen and Meyer [1966], Chandler et al. [1989]), including the GLIMPCE wide-angle array studies, which have imaged sub-crustal structures [Behrendt et al., 1990]. Seismic studies reveal an upper crustal rift basin, a high-velocity lower crust, and an overall thickened crustal section [Trehu et al., 1991; Hinze et al., 1992]. The interpretation is that the original crust was thinned to about half of its original thickness due to rifting, but the basin was filled in with extrusive volcanics and basin sediments, and the lower crust thickened with volcanic underplating, resulting in an over-thickened crust [Allen et al., 1995; Hinze et al., 1997].

An example of an Iowan trans-MCR profile is shown in Figure 9 [Chandler et al., 1989]. A question remains as to what underlies the region at depths of 50 - 300 km. The MCR formed concurrently with the accretion of the Grenville Terrane, and while there is a precedent for continental rifting due to subduction-generated plate suction forces [e.g., Storey et al., 1992], this situation is more complicated because (1) the MCR was simultaneously parallel to (Iowa) and perpendicular to (Michigan) the Grenville front, and (2) the MCR was soon after subjected to compressional stresses that stopped rifting and caused additional basin thickening. As a result, a hotspot is thought to supply the volcanism and crustal weakening [Hauser, 1996]. We aim to learn about the mantle flow that accompanied the creation of the MCR by determining the cross-MCR velocities at sub-crustal and sub-lithospheric depths. Active rifting can be associated with deep (> 500 km) slow seismic velocities, as under the Red Sea [Zhang and Tanimoto, 1992]. Is any deep thermal signature still present after 1.1 Ga? We also expect high-velocity depleted peridotites beneath the rift, and the depth of such anomalies may indicate the thermal state of the region 1 Ga ago.

MCR Seismicity. In general, there has been little seismicity along the MCR, with no obvious connection to MCR structures [Mooney and Morey, 1981; Mooney and Walton, 1986; Chandler and Morey, 1989; Hinze et al, 1992]. However, investigations of the MCR in Kansas have found seismicity associated with transform faults that truncate MCR segments [Steeples et al., 1989]. The Belle Plaine fault is a large MCR transform that extends from Minnesota into Iowa, and we plan to locate small events recorded at the Iowa stations to see if any seismicity is associated with this major fossil fault.

Key Questions:

  • What is the depth/extent of depleted ultramafic rocks beneath the MCR?
  • What is the depth/extent of any remaining deep thermal anomaly beneath the MCR?

3. Upper Mantle Anisotropy.

Determining the magnitude and direction of anisotropy in crust and upper mantle velocities beneath the FLED stations would serve two important objectives: (1) assess the effects of upper mantle anisotropy on the other proposed seismic analyses (in our studies of MOMA Sdiff and Pdiff slownesses and Sdiff splitting, we first corrected these phases for upper mantle anisotropy as reflected in SKS, PKS, and other core phases), and (2) resolve rock fabrics and tectonically induced strain [McKenzie, 1979; Nicolas and Christensen, 1987; Mainprice and Silver, 1993; Ribe, 1992; Wenk et al., 1991; Zhang and Karato, 1995].

A key question is how the fabric of the North American lithosphere was altered by the rifting that produced the MCR. Strain-induced fabrics within the lithosphere would likely endure because this region has seen relatively little deformation since the rifting occurred. We plan to measure shear-wave splitting in the abundant core phases (SKS, PKS, sSKS, etc.) in the 90\260 - 140\260 range from the seismogenic regions of Tonga, Taiwan, the Philippines, Indonesia and parts of Japan. As with MOMA data [Fouch et al., 1998b], we plan to analyze splitting in individual phases for a dependence of splitting on back-azimuth or incidence angles, diagnostic of multiple or dipping layers of anisotropy, and will use the stacking technique of Wolfe and Silver [1998] to determine average station splitting. We will look for evidence of anomalous MCR lithospheric fabric by comparing splitting at stations within and outside the rift zone (Figure 3). If the rifting fabric is parallel to extension and roughly normal to the MCR axis, as observed across a fast-spreading segment of the East Pacific Rise [Wolfe and Solomon, 1998], this signal can be clearly resolved because the fast direction of anisotropy would be nearly perpendicular to the ambient NE-ENE fast directions observed elsewhere in the mid-continental U.S. (Figure 1). If the rifting instead left a flow fabric whose fast direction is sub-parallel to the rift axis, as found in several continental rift zones [Sandvol et al., 1992; Gao et al., 1997], the rifting fabric would be revealed primarily by variations in splitting times.

The 25 new splitting parameters would also place additional constraints on broader models of anisotropy, lithospheric deformation, and deeper mantle flow in North America. In some recent models, upper-mantle anisotropy is attributed largely or exclusively to lithospheric deformation that is coherent from the crust to depths of 200 km or more [Helffrich et al., 1994; Silver, 1996], while in others observed anisotropy is attributed to shearing in the asthenosphere produced by drag from the lithosphere above [Vinnik et al., 1995b] or by small-scale convection [Makeyeva et al., 1992]. In Fouch et al. [1998b] we modeled shear-wave splitting measurements from MOMA and other stations (Figure 1), and found that while much of this signal could be matched by strain-induced anisotropy in mantle flow around and below a translating North American continental keel, anisotropy was still required in the lithosphere in a few localized regions. Splitting at the FLED stations would add considerable new resolution of splitting both in the keel interior, and across and outside the keel margin in the southeastern U.S.

Key questions:

  • Did Proterozoic rifting in the MCR substantially deform the lithosphere?
  • If so, what is the geometry and extent of this anomalous fabric?
  • How can it constrain the rifting process?
  • Is splitting best explained by lithospheric deformation, mantle flow around the keel, or both?

VI. Deployment

We plan to deploy 25 broadband stations during a 15-month period spanning June, 2001, to August, 2002. Data will be recorded at 40 sps on 2 or 4 Gb field disks. Streckheisen STS-2's or Guralp 3T's provide the long-period response needed for surface waves. One STS-2 will be provided by Washington University, and 24 will be borrowed from IRIS PASSCAL. As many sites as possible will be constructed with continuous power, as power failures have been the primary cause of lost data in our previous experiments.

A. Site Selection and Construction. Many sites have been targeted at locations corresponding to a college, university, or state park, where assistance can be obtained with site selection [Tyrrell Museum of Paleontology (AB), Theodore Roosevelt Natl. Park (ND), Drake (IA), Truman State (MO), Southeast Missouri State (MO), Meriwether Lewis Park (TN), Alabama A&M (AL), Columbus State (GA), Georgia Southwestern College (GA), Valdosta State (GA), Univ. of Florida (FL)]. These resources can provide help with station maintenance and protection. Private land owners, like farmers, have also been good site hosts, providing appropriate station protection, and will comprise the bulk of the northern sites. We have plenty of time to make these contacts, as the broadband PASSCAL sensors are not available until Spring, 2001 (J. Fowler, personal communication). Ideal sites are within a couple hundred meters of a continuous power source, sufficiently sloped for drainage, protected from vandalism, away from regular human or livestock traffic, at least 1 km away from major roads or railroads, away from trees and other vertical structures that can catch wind and block GPS satellite visibility, and have about 0.5 m of soil over bedrock (a bit deeper for Guralps). We are willing to move within a 10 km circle to get a good site (less for the MCR sites). Soil thickness maps are readily available from most State Agriculture Departments. In some regions bedrock is not available, and a soil installation is inevitable. These sites still provide excellent vertical components, but horizontal components contain long-period (T > 40 s) noise. Holes are dug with a portable-generator-powered jackhammer, and concrete pads poured in direct contact with bedrock. Vaults are constructed out of plastic cylinders that are about 0.5 m across, and seal tightly at the top. High-density styrofoam insulation is placed around the sensor, and the vault is buried to provide thermal insulation. Plastic "Tuff" boxes are used to store the two marine batteries, power board, DAS (with internal GPS clock preferred), field disk, breakout box and extra cable. This is also partially buried to provide insulation and reduce wind noise.

B. Sensor Deployment. Sensor deployment is done after the vaults are constructed because the concrete requires at least 1-2 days to cure, and doing everything in one trip results in either waiting around or constant back-tracking. Current magnetic-north directions are obtained from the USGS for all sites to aid in proper azimuthal orientation. Romex cable is shallowly buried for continuous power, though we estimate that about 5 sites will require site-generated solar power. Solar panels (4 minimum) are installed as far away as possible from the sensors. Trenches for water drains out of the bottom of the vaults are installed at this time. We predict that ∼5 days of power-trencher rental will be required for regions with tough soil or little topography. Sites requiring solar panels or power trenching are determined during the Vault construction.

C. Site Visits for Data Retrieval. Stations will be visited every 3-4 months, to inspect station performance (a 2 Gb disk will actually hold about 6 months of 3-channel data at 40 sps). We plan to carry out the site visits in 4 episodes: (1) day trips to Missouri stations, often used as training runs for students, (2) Tennessee-Florida stations, (3) Iowa stations, and (4) S. Dakota-Alberta stations. Field disks will be down-loaded to Jazz disks in the field or during motel evenings (tape malfunction has been a frustrating source of lost data). Data will be incorporated into the PASSCAL database on a dedicated portable Linux PC.

D. SEED Volume Production. We plan to produce quality-controlled, time-corrected SEED volumes of the data at Washington University shortly following site visits. MOMA was the first experiment where the PASSCAL database was installed at the host site. This was a dubious distinction, as recurrent bugs in the database delayed the final production of time-corrected data. However, the database is now working well, and we are using it successfully for the SEPA PASSCAL experiment [Wiens et al., 1994].

E. Schedule. 7/1999 - 3/2001 (Establish likely sites; Contact local Geology Departments, State Parks, etc.); 4/2001 - 5/2001 (Trip for final site selections and vault construction); 6/2001 (Sensor deployment); 8/2001, 12/2001, 4/2002 (Data retrieval - data processing and analysis); 8/2002 (Station removal).


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