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Magnetostratigraphy, paleomagnetic correlation, and deformation of Pleistocene deposits

Geologic Setting
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Magnetostratigraphy, paleomagnetic correlation, and deformation of Pleistocene deposits in the south-central Puget Lowland, Washington

Hagstrum, J.T.1, Booth, D.B.,2, Troost, K.G.,2Blakely, R.J.1


Paleomagnetic results from Pleistocene sedimentary deposits in the central Puget Lowland indicate that the region has experienced widespread deformation within the last 780 kyr. Three oriented samples were collected from unaltered fine-grained sediments mostly at sea level to determine the magnetostratigraphy at 83 sites. Of these, 47 have normal, 18 have reversed, and 18 have transitional (8 localities) polarities. Records of reversed- to normal-polarity transitions of the geomagnetic field were found in thick sections of silt near the eastern end of the Tacoma Narrows Bridge, and again at Wingehaven Park near the northern tip of Vashon Island. The transitional horizons, probably related to the Bruhnes-Matuyama reversal, apparently fall between previously dated Pleistocene sediments at the Puyallup Valley type section (all reversed-polarity) to the south and the Whidbey Island type section (all normal-polarity) to the north. The samples in general are of sufficient quality to record paleosecular variation (PSV) of the geomagnetic field, and a statistical technique is used to correlate horizons with significant agreement in their paleomagnetic directions. Our data are consistent with the broad structures of the Seattle uplift inferred at depth from seismic-reflection, gravity, and aeromagnetic profiles, but the magnitude of vertical adjustments is greatly subdued in the Pleistocene deposits.


Map of central Puget Lowland
Figure 1. Map of central Puget Lowland showing paleomagnetic sampling sites (circles). Filled circles indicate sites with normal-polarity remanent magnetizations, open circles indicate sites with reversed-polarity magnetizations, and half-filled circles indicate sites with transitional-polarity directions (Table 1). Inset at the lower right shows the location of the central Puget Lowland in Washington State. Click to enlarge image

The heavily populated Puget Lowland region (Fig. 1) has been the subject of an increasing number of scientific investigations designed to clarify the nature of its seismic hazards. The Lowland has had a large number of historical earthquakes relative to its surrounding areas, and most of the larger events have been within the subducting Juan de Fuca plate [Ludwin et al., 1991; Rogers et al., 1996]. Recent geologic investigations, however, have documented major prehistoric earthquakes in the overriding North American plate, in particular along the Seattle fault [Bucknam et al., 1992; Nelson et al., 1999]. Considering the potential consequences of a large earthquake in the modern Puget Lowland, relatively little information is available concerning the region's basic structural setting and paleoseismicity. Understanding the stratigraphy, structure, and deformation of Quaternary sedimentary deposits within the Puget Lowland is important for an accurate assessment of the region's seismic hazards.

Geologic mapping in the mostly unconsolidated sedimentary cover of the Puget Lowland is difficult primarily due to similar appearances of various Pleistocene glacial and nonglacial deposits mantling the area, and to poor exposure caused by abundant landslide deposits, dense vegetation, and urban development. We report here on a magnetostratigraphic study that was initially undertaken to provide a rudimentary understanding of the regional Pleistocene stratigraphy: reversed(R)-polarity paleomagnetic directions are assumed to indicate an age greater than ~780 ka, and normal(N)-polarity directions an age less than ~780 ka, the Bruhnes-Matuyama boundary.

The paleomagnetic data are apparently of sufficient quality, however, that paleosecular variation (PSV) of the geomagnetic field was recorded. Significant parts of a polarity reversal, probably the Bruhnes-Matuyama transition, are recorded at localities near the eastern end of The Tacoma Narrows Bridge and at Wingehaven Park near the northeastern end of Vashon Island (Fig. 1). Transitional directions are found at a number of other sites and potentially provide a high-precision stratigraphic marker horizon. Furthermore, similar PSV directions for sites of N, transitional, and R polarity have been correlated using a statistical technique developed by Bogue and Coe [1981]. Finally, we compare our paleomagnetic results with structural models developed using seismic reflection profiles beneath Puget Sound [Pratt et al., 1997] and tomographic data from the 1998 Seismic Hazards Investigation in Puget Sound (SHIPS; Brocher et al. [2001]).

Geologic Setting

The oceanic Juan de Fuca plate is the northernmost remnant of the Farallon plate subducting beneath North America, and its oblique convergence with the continental margin is the source of great subduction-zone earthquakes [Atwater and Hemphill-Haley, 1997], complex upper-plate folding and faulting [Johnson et al., 1999], and Cascade-arc volcanism [Smith, 1993]. Based on Neogene deformation, paleomagnetic rotations, and geodetic data, the Cascadia forearc appears to be migrating northward and breaking up into large rotating blocks associated with dextral transpression [Wells et al., 1998]. The Puget Lowland lies within this transpressional zone in northwestern Washington, overlying a major crustal boundary between Eocene basement rocks of the western Coast Range province and pre-Tertiary rocks of the eastern Cascade province [Wells and Coe, 1985; Johnson et al., 1996].

Seismic reflection profiles beneath Puget Sound indicate subhorizontal Paleogene and Neogene sedimentary rocks deformed by west- to northwest-trending faults and folds [Pratt et al., 1997]. The stratigraphy in the Lowland consists of the Crescent Formation basalts overlain in the Seattle basin by upper Eocene marine sedimentary strata, shallow-marine turbidites of the upper Eocene to Oligocene Blakeley Formation, non-marine Miocene sediments of the Blakely Harbor Formation, and by Quaternary deposits. Asymmetric subsidence of the Seattle basin indicates continued contractional deformation since Eocene time [Johnson et al., 1994].

A primary subsurface feature is the Seattle uplift, with south dipping (~20°) bedrock strata on its south flank, and steeply north dipping (50°-90°) strata and the east-trending Seattle fault on its north flank. Pratt et al. [1997] interpreted this and other uplifts as fault-bend and fault-propagation folds, and suggested that the Puget Lowland is underlain by a large north-directed thrust sheet (thin-skinned deformation) bounded by faults along the Cascade and Olympic Ranges to the east and west [Johnson et al., 1994], respectively. Brocher et al. [2001], on the other hand, interpret the Seattle uplift as a pop-up structure [Wells and Weaver, 1993] bounded by the steeply dipping Seattle fault to the north and Tacoma fault to the south. The Tacoma fault is inferred from gravity and magnetic data and a seismic-velocity gradient similar to that associated with the Seattle fault. In addition, the Tacoma fault is inferred from the magnitude of structural relief along this zone, particularly to the west. Conversely, structural relief along the Seattle fault zone decreases to the west, and Brocher et al. [2001] suggest that this relationship likely results from the transfer of strain between the Seattle and Tacoma faults.

Conceptual Quaternary stratigraphic framework Figure 2. Conceptual Quaternary stratigraphic framework for the Puget Lowland showing the younger Whidbey Island type section [Easterbrook et al., 1967] with the addition of the Olympia beds [Troost, 1999] and the older Puyallap Valley type section [Crandell et al., 1958]. At present, an age data gap exists between the older reversed-polarity section (›780 ka BP) and the younger normal-polarity section (‹250 ka BP). Modified from Troost [1999]. Click to enlarge image

The glacial and interglacial deposits overlying Tertiary basement rocks in the southern Puget Lowland (Fig. 2) indicate that the area was glaciated at least six times during the Pleistocene Epoch [Easterbrook, 1994]. Near Tacoma, these deposits are approximately 400 m thick [Jones, 1996]. Correlations between units have been difficult to make due to the lack of distinctive lithologic or textural features and to problems in dating materials of this composition and age. More recently, laser-argon, fission-track, thermolumines-cence, amino acid, and paleomagnetic techniques have been employed to establish a chronology for these Pleistocene sediments, particularly north and east of the Seattle-Tacoma area [Blunt et al., 1987; Easterbrook, 1994; Troost, 1999; Troost and Booth, 1999].

The last glacial advance in the Puget Lowland was the Puget lobe of the Cordilleran Ice Sheet, culminating ~15 ka ago during the Vashon stade of the Fraser glaciation [Booth, 1987]. In the Puyallup River valley, one of the type sections of older, pre-Vashon glacial and nonglacial deposits is exposed (Fig. 2). It consists of drift and outwash of the Orting, Stuck, and Salmon Springs glaciations [Crandell et al., 1958]. All of these glacial materials and their interbedded nonglacial deposits have R-polarity remanent magnetizations and ages greater than ~780 ka [Easterbrook et al., 1988; Easterbrook, 1994]. To the north on Whidbey Island (~50 km north of Seattle), the upper type section of Pleistocene deposits (Fig. 2) consists of the Double Bluff, Possession, and Vashon glacial deposits and intervening glacial and nonglacial deposits [Easterbrook et al., 1967]. Here, the entire section has N-polarity magnetizations and ages less than ~780 ka [Easterbrook, 1994]. No deposits between the Double Bluff and Salmon Springs glaciations (Fig. 2) have yet been identified in the Puget Lowland [Richmond and Fullerton, 1986]. Mapping efforts in the Tacoma area have identified the Vashon glacial drift and at least two older drifts of N polarity (this study) which may fall between the Double Bluff and Salmon Springs deposits based on preliminary luminescence dates [Troost, 1999; Mahan et al., 2000].


Generally, three oriented samples were collected from unweathered (dark gray), fine-grained (silt) glacial and interglacial lake deposits (Fig. 2) at each of 86 sites in the central Puget Lowland. Horizontal benches were excavated and leveled with a circular bubble level, and then vertical pedestals were carved over which plastic sample boxes (volume = 6 cm3) were fitted. Before removal, one top edge of each box was marked and oriented using a magnetic compass. Local bedding was also recorded to correct for deformation since original deposition. The sites are mostly at sea level in beach exposures around Puget Sound, but some were also collected in road-cut (N3-N6, N16, N17, R12, N27, N28, N37, N38, N40), stream-bank (N1, N2, R16, N39, N44), and hillside (R3, T1, T3-T6, T8, N35, N46) exposures (Fig. 1; Table 1).

Orthogonal projection of alternating-field  demagnetization vector endpoints Figure 3. (a) Orthogonal projection of alternating-field (AF) demagnetization vector endpoints for a sample from site N10 showing a univectorial normal-polarity remanent magnetization. (b) Vector plot of AF demagnetization data for a sample from site T9 showing a univectorial transitional-polarity magnetization. Filled symbols in both plots indicate projections onto the horizontal plane, and open symbols, onto the vertical plane. (c) Equal-area stereoplot showing intersecting remagnetization circles fitted to demagnetization data for samples from site R10. Dots indicate poles to the great circles projected from the lower hemisphere, and the open circle indicates the mean direction (reversed-polarity) for this site projected from the upper hemisphere. The intersection is tightly constrained, and the 95% confidence ellipse is correspondingly small. Click to enlarge image

Step-wise alternating field demagnetization indicates that these sediments carry stable characteristic remanent magnetizations (Fig. 3a, 3b). Three sites with unstable magnetizations were discarded from further analysis. The normal-polarity characteristic magnetization vectors were usually isolated between 20 and 100 mT. Least-squares lines are fitted to the sample demagnetization data [Kirschvink, 1980], and Fisher [1953] statistics are calculated for single-component site-mean directions and for the overall mean directions. For R-polarity samples, N-polarity components were first removed or, more often, the N-polarity overprint was removed simultaneously with the characteristic R-polarity component. Converging demagnetization planes and Bingham statistics [Onstott, 1980], therefore, were used to determine the characteristic magnetization direction and error limits, respectively, for most of the R-polarity sites Fig. 3c.

Equal-area stereoplots
Figure 4. Equal-area stereoplots of normal-polarity, reversed-polarity, and transitional-polarity site-mean directions with their 95% confidence limits. Numbers are keyed to site data listed in Table 1. Filled circles are projected from the lower hemisphere, and open circles are projected from the upper hemisphere. Small triangles indicate site-mean directions with too few samples to calculate meaningful statistics (Table 1). Click to enlarge image

In this procedure, great circles were fitted to the demagnetization end-point data that are curved rather than linear in vector plots due to overlapping coercivities for the two components of remanent magnetization. The great circles are expected to intersect at the component's direction having the higher coercivity range. The technique works best if the lower-coercivity components have random directions, making the statistical certainty of the characteristic direction greater and its 95% confidence limits correspondingly smaller and more circular. If both components are non-random, however, the intersection point of the great circles might be nearer the higher-coercivity direction or the antipode of the lower-coercivity direction, depending on how well the directions are represented by the sample population. In such cases, the error ellipses are more elongate (Fig. 4) and the mean directions are less reliable. Samples from 47 sites have N-polarity mean directions, 18 sites have R-polarity directions, and 18 sites (8 localities) have transitional directions (Fig. 4, Table 1). Two R-polarity sites (R3, R4) have an associated fission-track age of ~1.1 Ma [Booth et al., in press]. An optically-stimulated luminescence (OSL) date for sediments just east of site R4 indicate an age of ~250 ka [Mahan et al., 2000]. OSL and thermal luminescence (TL) dates for sands and silts at Point Defiance (N11, N12, N19) and at the intersection of I-5 and Atlantic Street (N27, N28) indicate ages between 200 and 300 ka and of ~70 ka, respectively. OSL dates at Garfield Park in Tacoma (N17) and near Dash Point (N25) indicate ages ›107 ka and of ~180 ka, respectively [Mahan et al., 2000].

Equal-area stereoplot Figure 5. Equal-area stereoplot of site-mean directions for (a) a thick silt section just north of the eastern abutment of The Tacoma Narrows Bridge, and (b) the section at Wingehaven Park (Fig. 1). Sites R12 and R18 include the stratigraphically lowest samples at each site and have reversed-polarity mean directions. Site N37 is the stratigraphically highest site at Wingehaven Park and has a normal-polarity mean direction. Stratigraphic distances are indicated between sites as the paleomagnetic direction swings between reversed polarity and normal polarity; the path of the transition at Tacoma Narrows is shown by the dashed line. For comparison, that same path is superimposed on the Wingehaven Park data (dotted line), along with the apparent continuation of the transition to fully normal polaity (solid line). Inverted triangles indicate the normal and reversed directions of the present-day geomagnetic field in the central Puget Lowland. Click to enlarge image

A 20 m-thick silt section near the eastern end of The Tacoma Narrows Bridge preserves part of a R-to-N transition. Similarly, at Wingehaven Park on Vashon Island samples were collected above, below, and within a R-to-N polarity transition (Fig. 5). Directions inferred as transitional were also found along the western shore of Tacoma Narrows (T13), along Puget Sound's eastern shore (T8, T11), and at the northern tip (T7) and along the western shore of Vashon Island (T2, T18; Fig. 1).

The site-mean statistics are often remarkably good for only three sample directions, although the overall dispersion of VGPs for both polarity groups (SN+R = 30°) is significantly greater than a model value for the full range of secular variation at this latitude (SF = 17° ± 1°; McFadden and McElhinny [1984]). Although the higher observed dispersion could have been caused by inaccuracies in the sediment recording process (e.g., bottom paleocurrents), differential vertical-axis rotations, unrecognized or incorrect stratal tilts, unrecognized transitional directions, and/or unrecognized overprinting, the excess dispersion can also be attributed to higher within-site dispersion due to the low number of samples per site. The near-antipodal mean directions of the N- and R-polarity groups, however, indicate that PSV and other sources of error have been averaged out.

Assuming that the silt deposits at sites with well-grouped sample directions (a95 ‹15°) reliably record PSV, the observed paleomagnetic directions can be statistically compared with one another to estimate the relative likelihood of sites having similar directions due to coincidence or to significant agreement. Bogue and Coe [1981] initially developed a statistical method to correlate paleomagnetic directions from individual Columbia River basalt flows. Their method is based on the observation that the geomagnetic field direction at any given locality tends to be near the expected dipole field direction. Thus, two similar but unusual directions away from the expected field direction are more likely acquired simultaneously than two similar directions near the expected direction. This technique is most accurate when applied over short periods of time relative to the frequency of PSV. In this study, the period of time over which the sampled sediments were deposited is relatively long, so correlations indicated by the statistical comparison are less certain and therefore only the strongest correlations are considered. In addition, the overall mean direction (Table 1) was substituted for the expected dipole field direction because the observed mean has a shallower inclination and a slightly more counterclockwise declination.

In Bogue and Coe's [1981] method, two hypotheses are tested. The ‘random' hypothesis (Hr) holds that similar paleomagnetic directions are random samplings of the geomagnetic field, and the ‘simultaneous' hypothesis (Hs) holds that the directions were acquired under the same geomagnetic field. The calculated probabilities (P) that similar paleomagnetic data (D) have arisen from either hypothesis (P(D:Hr) or P(D:HS)) are given in Table 2. P(D:Hr) is calculated using a spherical distribution model based on Fisher's [1953] probability function, and P(D:HS) is the significance level from McFadden and Lowes's [1981] test of the null hypothesis stating that the two sample means are from populations having the same mean but different k (concentration) values. The ratios of P(D:HS)/P(D:Hr), also given in Table 2, indicate the relative likelihood of Hs versus Hr.


The fine-grained deposits of the Puget Lowland apparently record polarity and PSV of the geomagnetic field. As previously mentioned, however, a number of error sources might have affected the accuracy with which the fine-graineddeposits recorded the ambient geomagnetic field. Paleocurrent alignment of magnetic grains is unlikely in lake bottom environments, and lake sediments have provided consistent and reproducible records of PSV in North America [Lund, 1996]. Vertical-axis rotations are also unlikely in deposits this young, and although unrecognized stratal tilts could contribute to the error they too would be relatively minor. The division between transitional directions and extreme PSV directions is arbitrary, and some of the highly dispersed N-polarity directions in Figure 4 might actually be transitional directions (see below). Unrecognized overprinted N-polarity directions are also unlikely because uniform unaltered sediments were sampled in which R-polarity and transitional directions were also found. Increased dispersion due to the low number of samples collected could not be avoided without greatly increasing the sampling time.

The overall average inclinations for the fine-graineddeposits (Table 1; Fig. 4) are also too shallow compared to the expected dipole-field direction. The shallowing of inclinations is most likely due to compaction of the fine-grained lake sediments [Anson and Kodama, 1987], particularly in the older Quaternary deposits due to loading by ice sheets during the multiple Pleistocene glaciations. Because the degree of compaction is probably crudely similar at equivalent stratigraphic levels, comparisons of paleomagnetic directions to determine equivalent horizons are most likely valid. The transitional paleomagnetic directions for sites collected at Tacoma Narrows and at Wingehaven Park are plotted in Figure 5. Both of the stratigraphically lowest sites at Tacoma Narrows and Wingehaven Park have R polarity (R12, R18), and the highest site at Wingehaven Park has N polarity (N37). Intermediate directions at both localities are transitional, and clearly a R-to-N reversal is recorded in the fine-grainedsediments. Directions at other sites with northerly declinations were considered transitional if their mean inclinations were ‹25°. The cut off value is arbitrary, however, and shallow-inclination N-polarity directions, such as those for sites N9, N15, and N27, might also be transitional directions.

Although the transitional directions could be associated with a number of subchrons within the Bruhnes polarity chron [Champion et al., 1988], it is most likely related to the Bruhnes-Matuyama transition at ~780 ka. A fission-track date on an interbedded tephra layer near site R4 indicates that Matuyama-aged sediments have been sampled there. A finite 14C date of 44,880 ± 3050 ka near and stratigraphically above the eastern Tacoma Narrows locality [Troost, 1999] preclude these transitional directions from being related to the Laschamp event (40 ka). Furthermore, the OSL and TL dates nearby and at Point Defiance (N11, N12, N19) between 200 and 300 ka indicate that the N-polarity sediments are below either the well-defined Jamaica (~180 ka) or Blake (~110 ka) events. Older subchrons have been proposed within the Bruhnes chon, but are not as well established.

Paleomagnetic correlations of the transitional directions also indicate that parts of the same geomagnetic reversal were sampled across the study area. The calculated probabilities and relative probabilities of the random (Hr) or simultaneous (Hs) hypotheses [Bogue and Coe, 1981] are given in Table 2. The first normal-polarity comparison in Table 2 is between site N1 (Fig. 4) and 9 other sites with similar directions. All of these directions are near the overall mean direction (d < 20°), and their relative probability values are correspondingly low (‹40) and mixed between favoring Hr or Hs. In contrast, a correlation test between sites N9 and N15 indicates that Hs is more likely than Hr by a factor ›9999, and so these two sites are probably within the same magnetostratigraphic horizon. Transitional directions at the east shore of Tacoma Narrows (T4, T5) are identical to those at Wingehaven Park (T14, T15), and that at Seahurst Park (T11) is widely correlated with directions at Redondo (T8), Peter Point on Vashon Island (T10), and on the west shore of Tacoma Narrows (T13). In addition, some of the northernmost N-polarity directions (N31, N33) are correlated with some of the southernmost sites (N16, N38).

Sampling sites and magnetic polarities Figure 6. Map of the central Puget Lowland showing sampling sites and magnetic polarities (as in Fig. 1) superimposed on an aeromagnetic map for the region [Blakely et al., 1999]. Bold dashed lines (B and C) indicate the margins of the Seattle uplift and the limit (A) of the southward dipping ramp at the southern edge of the uplift (after Pratt et al. [1997]). The dotted line north of the Seattle uplift marks the trace of the Seattle fault. Fine dashed lines connect sites that have been statistically correlated based on their atypical, but similar, paleomagnetic directions (Table 2). Click to enlarge image

In Figure 6, the distribution of N-, transitional-, and R-polarity sites are shown superimposed on an aeromagnetic map of the central Puget Lowland [Blakely et al., 1999]. Also shown are most of the calculated correlations between sites having relative probability factors for Hs of several thousand or more (Table 2). The aeromagnetic data show the local influence of human activity, such as The Tacoma Narrows Bridge (near T1 and N46) and at the Tacoma harbor docks (near R4). In addition, the pattern of N- and R-polarity sites near the southern end of Vashon Island appear to sharply define the southern limit of the Seattle uplift at its boundary with the adjacent Tacoma Basin (NW to SE line B). This boundary is also well defined by gravity and seismic tomography data [Brocher et al., 2001].

The pattern of paleomagnetic polarity Fig. 6 appears unrelated to the pattern of aeromagnetic anomalies indicating that the anomalies are not due to differential uplift of strongly magnetized basement rocks such as the volcanic Crescent Formation. The polarity pattern is the result of the elevation of sites, past erosion, and tectonic movements. An estimate of the elevation of each site is given in Table 1, and clearly plays a role in polarity at the thick sections with transitional directions sampled at The Tacoma Narrows Bridge and Wingehaven Park. The N-polarity sites at Christiansen Road (N3-N6) on the western coast of Vashon Island, just across from R-polarity sites on beaches of the eastern Kitsap Peninsula (Fig. 1), have much higher elevations (20-50 m). Moreover, N-polarity sites are always found above R-polarity sites indicating that R-polarity subchrons within the Bruhnes chron have not been sampled.

Deposition of fine-grainedsediments over a surface with erosional relief might also account for the close proximity of N- and R-polarity sites, but is difficult to evaluate because of limited exposures in the region. Tectonic uplift can cause the exposure of R-polarity sediments, and the majority of sites in such sediments over the Seattle uplift Fig. 6 indicate that uplift has continued since ~780 ka. The abrupt change in polarity across the southern boundary of the Seattle uplift might also indicate that the Tacoma fault intersects the land surface. In addition, a dip-slip fault along the Tacoma Narrows is indicated by the down dropping of sites T13 and N46 on the west relative to sites R12, T1, and T3-T6 on the east.

Sampling sites and magnetic polarities Figure 7. North to south magnetic [Blakely et al., 1999] and gravity [Brocher et al., 2001] profiles (line P-P' in Fig. 6) across the south-central Puget Lowland are shown in (a) and (b), respectively. A model is shown in (c) in which seismic tomography data (depth to 4.5 km/sec velocity rocks) is used to constrain the gravity fit, and reversed-polarity slabs of Crescent Formation are used to fit the magnetic profile. The model is interpreted in (d) with the Seattle fault as a south-dipping reverse fault and the southern margin of the Seattle uplift as possibly a south-dipping ramp. Thickness of the thin layer of Quaternary sediments in (c) and (d) is based on data from Jones [1996]. Click to enlarge image

To determine the source of the aeromagnetic anomalies, gravity and magnetic profiles along line P to P' (Fig. 1) were fitted by a simple subsurface model shown in Figure 7. The fit of the gravity profile is constrained by SHIPS tomographic data indicating the depth to rocks having 4.5 km/sec velocities [Brocher et al., 2001], which is presumably the top of the Crescent Formation (Fig. 7c). Not surprisingly, the fit to the gravity profile is excellent, but the concurrent fit to the magnetic profile (not shown), assuming a uniformly magnetized Crescent Formation, is poor. Slabs of reversely magnetized rock must be added to the model so that the calculated profile matches the observed magnetic profile (Fig. 7a, 7c). Although the Crescent Formation is reversely magnetized at the surface in some places [Beck and Engebretson, 1982; Globerman et al., 1982; Wells and Coe, 1985], locations of reversely magnetized rocks in Figure 7c are constrained only by the shape of the aeromagnetic data. In this model, the Seattle fault is interpreted as a south-dipping reverse fault and the southern margin of the Seattle uplift as a south-dipping ramp (Pratt et al. [1997]; Fig. 7d). The reversely magnetized layers within the Crescent rocks also dip southward and might reflect stratigraphy within the formation.


In this study paleomagnetic directions from fine-grained unaltered glacial and interglacial deposits have been used to define the area's magnetostratigraphy (Figs. 1, 6). Remarkably, the horizontal plane of sampled exposures (~sea level) intersects a R-to-N polarity transition that is most likely related to the Bruhnes-Matuyama geomagnetic reversal (Fig. 5). Vertical adjustments on the order of ‹10 m could determine whether a sediment with N, R, or transitional directions was sampled. The transitional horizon is at least 2 m thick and serves well as a stratigraphic marker horizon within the previously undated sediments between the Salmon Springs and Double Bluff glacial deposits (Fig. 2). Seismic-reflection data analyzed by Pratt et al. [1997] image subsurface structure to depths of several km, whereas seismic tomography [Brocher et al., 2001] does so to depths of 25 to 30 km.

Equal-area stereoplots
Figure 8. North-to-south cross sections beneath the central Puget Lowland based on (a) seismic reflection profiles analyzed by Pratt et al. [1997] and on (b) the 3-dimensional seismic velocity model of Brocher et al. [2001]. Dots indicate hypocenters of local earthquakes projected E-W onto the cross sections. In (a), a thin-skinned model is shown in which the Seattle fault is a thrust fault that shallows with depth and merges with a mid-crustal decollement. Light gray areas indicate Miocene and younger deposits, and darker gray areas indicated Eocene and Oligocene deposits (after Pratt et al. [1997]). In (b), a thick-skinned deformational model is shown in which the steeply dipping Seattle and Tacoma faults bound the Seattle uplift to the north and south, respectively. The steeply dipping faults connect at high angles with a lower crustal decollement at the base of the Crescent Formation. Focal mechanisms for the 1995 M = 5 Point Robinson and the 1997 M = 5 Bremerton earthquakes are also shown and are interpreted as having occurred on the Tacoma and Seattle faults, respectively (after Brocher et al. [2001]). Click to enlarge image

In Figure 8, cross-sections of the Puget Lowland are shown depicting the thin- and thick-skinned structural models for the region based on the reflection and tomographic techniques, respectively. Overall, the paleomagnetic data conform to the Seattle uplift: R-polarity sites are mostly found above this structural feature (Fig. 7). The sharp boundary between the N- and R-polarity data along the Seattle uplift's southern edge (line B; Fig. 6) is consistent with a fault structure (Tacoma fault?) that apparently reaches the surface. The paleomagnetic correlation of transitional and N-polarity sites in the southern part of the study area with sites in the northern part, however, implies less deformation of the Pleistocene sediments than of the underlying Tertiary deposits.

Regional deformation of the Puget Lowland is a result of the ongoing convergence of the Juan de Fuca and North American plates expressed through both faulting and folding. Active faults in the Puget Lowland have been inferred to offset Quaternary deposits, and although only the Seattle fault has had a clear history of late Holocene surface rupture [Nelson et al., 1999], future major earthquakes are certain to occur within the region. The location and nature of active faults, as well as the overall structure beneath the Puget Lowland, are at present still open questions and more work is needed to decipher the region's complex structural setting and paleoseismicity.

Acknowledgements. We thank S. Bogue, T. Brocher, P. Haeussler, S. Johnson, T. Walsh, and particularly R. Wells for helpful discussions; P. Haeussler for collecting samples at several localities; and T. Brocher for providing a preprint of their manuscript on the SHIPS data. We also acknowledge S. Bogue, P. Haeussler and S. Johnson for constructive reviews of the manuscript, and B. Graham and D.B. Bridges for assistance in the field and laboratory.


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Related Pacific Northwest Urban Corridor Geologic Mapping Project publications


Richard J. Blakely, U.S. Geological Survey, 345 Middlefield Road, Mail Stop 989, Menlo Park, California 94025. (e-mail:

Derek B. Booth, University of Washington, Box 352700, Seattle, WA 98195. (e-mail:

Jonathan T. Hagstrum, U.S. Geological Survey, 345 Middlefield Road, Mail Stop 937, Menlo Park, California 94025. (e-mail:

Kathy G. Troost, University of Washington, Box 351310, Seattle, WA 98195. (e-mail:


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