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Active Tectonics: Studies in Geophysics 6 Coastal Tectonics KENNETH R.LAJOIE U.S. Geological Survey, Menlo Park INTRODUCTION Between one-third and one-half of the Earth’s marine coastlines lie along or near tectonically and seismically active plate boundaries (Inman and Nordstrom, 1971). Many of the world’s major population centers lie along these active coastlines and, therefore, are vulnerable to adverse impacts of large earthquakes. Fortunately, there has been remarkable progress over the past two decades in the field of coastal tectonics, the study of recent crustal deformation and paleoseismicity in coastal regions. Of particular importance is progress in establishing earthquake recurrence, dating recent fault activity, and measuring rates of recent crustal deformation, all of which are crucial in determining earthquake potential and assessing seismic hazards and risk. Much of the recent progress in coastal tectonics stems directly from the development of new dating techniques that have led to a better understanding of Quaternary [past 2 million years (m.y.)] sea-level fluctuations and their relationship to marine strandlines (abandoned or relict marine shorelines), which are among the most numerous and widespread tectonic markers in the Quaternary geomorphic record (Figures 6.1, 6.2, and 6.3). Sea level is the common and unifying element of coastal tectonics. Present sea level, the universal datum for measuring elevation, is the most convenient reference for detecting ongoing vertical crustal movement and assessing short-term tectonic stability in coastal areas. Accordingly, coastal tectonics often includes analysis of tide-gauge data (Figure 6.4) and other historical and archeological information that might reveal apparent sea-level changes. Past sea levels derived from the geologic record comprise a composite datum for measuring long-term crustal movement. Consequently, coastal tectonics also includes mapping and dating marine strandlines and separating the vertical crustal movements from the past sea-level fluctuations they si- FIGURE 6.1 Emergent marine strandlines form steplike terraces on the Palos Verdes Peninsula of southern California. These and similar strandlines on technically active coastlines record both crustal uplift and major glacio-eustatic sea-level fluctuations (see Figure 6.6). The lowest strandline terrace is about 100 ka and the highest is about 1000 ka. The highest point on the peninsula is about 450 m above sea level. Modified from Davidson (1889).
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Active Tectonics: Studies in Geophysics multaneously record. Because marine strandlines are the physical records of past sea levels, the study of sea-level history is an integral part of coastal tectonics. Time-transgressive sequences of displaced and deformed Pleistocene [2 m.y. to 10,000 yr (10 ka) ago] marine strandlines document patterns of and, where dated, yield rates of continual, long-term crustal deformation (Figures 6.1 and 6.2). Sequences of emergent Holocene (past 10 ka) strandlines (Figure 6.3), which occur only along the most rapidly uplifting coastlines, commonly record abrupt coseismic uplift events of 1–15 m (see Figures 6.25–6.28 below) that cumulatively comprise the long-term deformation recorded by Pleistocene strandlines. Consequently, sequences of Holocene strandlines often record past earthquakes and, where dated, yield earthquake periodicity and provide a means of forecasting future seismic events. In many areas, apparent sea-level changes documented by tide-gauge records (Figure 6.4) or subtle shifts in the location of the modern shoreline reflect ongoing vertical crustal movement. Frequently, this movement is opposite in sense to long-term trends and, therefore, may represent postearthquake crustal relaxation or pre-earthquake strain accumulation. Coastal tectonics includes the study of both onshore (emergent) and offshore (submergent) marine strandlines and structural features. However, offshore coastal tectonics is a highly specialized field and is beyond the scope of this brief review, which focuses mainly on the formation and deformation of emergent marine strandlines and stresses their importance in determining the style and measuring the rates of recent crustal deformation, especially in highly active coastal regions. Most examples of strandline displacement and deformation cited in this review reflect sustained tectonic processes, but a few, included mainly for comparative purposes, reflect transitory volcanic and glacio-isostatic crustal deformation. COASTAL MORPHOLOGY AND TECTONIC SETTING On global and regional scales coastal morphology correlates closely with tectonic setting (Inman and Nordstrom, 1971). The greatest geomorphic contrast is between the subdued coastlines along passive continental margins and the rugged coastlines along convergent plate boundaries. Along most coastlines modern (active) coastal landforms are similar to their Pleistocene counterparts, which suggests that current tectonic and coastal processes have been fairly uniform over considerable periods of time. Most exposed coastlines along passive continental FIGURE 6.2 Cross-sectional profiles of emergent Pleistocene strandline terraces. A, Erosional strandlines near Santa Cruz, California. Solid line is existing topographic profile. Dashed line represents original profile of each strandline terrace, which consists of a relict wave-cut platform backed by a relict sea cliff. The intersection of the platform and the sea cliff is the shoreline angle, which closely approximates the paleoshoreline. Seaward thinning wedges of alluvial sediment derived from the degrading sea cliffs overlie the wave-cut platforms. A 120-ka strandline was removed from this section when the next lower (104 ka) platform was cut; the 120-ka strandline occurs a few kilometers north of this site and is projected onto this cross section. The three lowest strandlines were dated by paleontological, amino acid, and geomorphic techniques. The average uplift rate (0.35 m/ka) was derived from the vertical displacement (28 m) of the 82-ka strandline. The three highest strandlines were dated by extrapolation of this uplift rate and by numerical analysis of the progressively gentler slopes of successively higher (older) relict sea cliffs. Modified from Hanks et al. (1984). B, Depositional strandline terraces on the Huon Peninsula, Papua New Guinea. Each platform is a relict coral reef. U-series dates on fossil corals from these strandlines yield a history of glacio-eustatic fluctuations that serves as a tectonic datum for measuring vertical tectonic movements on other coastlines throughout the world (see Figure 6.6). The maximum average uplift rate of 4 m/ka on the Huon Peninsula was derived from the maximum vertical displacement (500 m) of the 120-ka strandline. Note that the three lowest strandlines on the Santa Cruz coastline correlate with the three highest strandlines in this sequence. Modified from Chappell (1974a).
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Active Tectonics: Studies in Geophysics FIGURE 6.3 Cross-sectional profiles of Holocene strandlines. A, Erosional terraces on the Boso Peninsula of Honshu, Japan. These strandlines, unlike their Pleistocene counterparts, represent periodic coseismic uplift events, not sea-level fluctuations (see Figures 6.7, 6.25, 6.26, 6.27, and 6.28). The lowest platform (IV) records uplift that accompanied two major earthquakes in 1703 and 1923, and the three highest strandlines were dated by radiocarbon analysis of fossil shells. The average uplift rate is 3.9 m/ka (see Figure 6.25). Modified from Matsuda et al. (1978). B, Depositional strandlines (beach ridges) at Te Araroa on North Island, New Zealand. Each ridge most likely represents a storm event not a coseismic uplift event. If uplift events are recorded in this sequence of strandlines, they are indistinguishable from storm events. The average uplift rate derived from the highest beach ridge (6 ka) is 1 m/ka. Modified from Garrick (1979). margins consist of broad coastal plains bordered offshore by wide continental shelves and gentle continental slopes. These relatively stable coastlines are characterized by depositional landforms such as broad sandy beaches and offshore barrier bars and at low latitudes by broad coral reefs. The southeast coast of North America and the northeast coast of Australia are typical passive-margin coastlines with low-to-moderate topographic relief and subdued depositional landforms. Much of the former is bordered offshore by barrier bars (Oaks and Du Bar, 1974), and most of the latter by extensive coral reefs (Hopley, 1983). Long-term tectonic stability along most passive-margin coastlines is expressed stratigraphically by undeformed continental and marine sediments that underlie flat coastal plains and continental shelves and geomorphically by broad accretionary strandline terraces that consist of subdued beach ridges separated by abandoned tidal flats (Oaks and Du Bar, 1974). However, rapid sediment accumulation, such as at the mouth of a large river, may isostatically depress an otherwise stable passive-margin coastline (Figure 6.4C) (Fisk and McFarlan, 1955; Hicks and Crosby, 1974). Also, occasional large earthquakes, such as the 1886 Charleston, South Carolina, seismic event on the Atlantic coast of North America (Hays and Gori, 1983), indicate that passive continental margins are not completely aseismic, even though there is little stratigraphic or geomorphic evidence in these areas of recent, near-surface crustal deformation. In contrast to the subdued coastlines along passive plate boundaries, most coastlines along or near active continental margins consist of coastal hills or mountains bordered offshore by narrow continental shelves and FIGURE 6.4 Tide-gauge records. A, San Francisco, California—the secular drift of this record (~2 mm/yr) is similar to that found in other parts of the world and, therefore, probably represents eustatic rise in sea level. B, Juneau, Alaska—the extreme apparent drop in sea-level represents crustal uplift due to either tectonic uplift or residual glacio-eustatic rebound. C, Mississippi delta—the apparent rise in sea-level represents subsidence due to sediment compaction and isostatic adjustments of the crust to the sediment load of the Mississippi delta. The eustatic rise (2 mm/yr) was subtracted from the relative rates to obtain the uplift and subsidence rates. Modified from Hicks and Crosby (1974).
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Active Tectonics: Studies in Geophysics deep submarine trenches, particularly in the circum-Pacific region. These tectonically active coastlines are characterized by rugged erosional landforms such as steep sea cliffs and rocky headlands, islands and sea stacks. However, at low latitudes coral reefs commonly form narrow depositional platforms even along rugged coastlines (Figure 6.2B)(Chappell, 1974a). The mountainous west coast of South America is an active-tectonic coastline that geomorphically reflects regional crustal uplift and subsidence related to rapid plate convergence and resultant subduction along the offshore Peru-Chile Trench (Plafker, 1972). The hilly-to-mountainous coast of California in western North America is an active-tectonic coastline that geomorphically reflects slower regional uplift and local basin subsidence related to oblique plate convergence and resultant large-scale, right-lateral displacement across the San Andreas Fault system. Long-term crustal instability along most active-tectonic coastlines is expressed stratigraphically by folded and faulted marine sediments that fill youthful structural basins and geomorphically by narrow uplifted and deformed Pleistocene strandline terraces that notch steep coastal slopes (Figures 6.1 and 6.2). Short-term instability along extremely active coastlines is expressed geomorphically by emergent Holocene strandline terraces (Figure 6.3) and by dramatic changes in the location and configuration of modern shorelines that result from rapid vertical crustal movements (Figure 6.4). Most of the Earth’s seismicity occurs along or near active-tectonic coastlines (Tarr, 1974), and many large earthquakes in these areas are recorded geomorphically by emergent Holocene strandline terraces (Figure 6.3A; also see Figures 6.25–6.28 below). The subdued coastlines bordering the shallow epicontinental seas in North America (Hudson Bay) and Fennoscandia (Bay of Bothnia) are exceptions to the general rule that coastal regions undergoing rapid crustal deformation are characterized by marked topographic relief. However, the rapid crustal uplift recorded so dramatically by sequences of highly emergent Holocene strandlines in these recently deglaciated areas reflects transitory isostatic rebound (see Figure 6.10 below) not sustained tectonic deformation. MARINE STRANDLINES Marine strandlines are the geological and historical records of former sea levels. In the geologic record marine strandlines are the depositional and erosional remains of abandoned marine shorelines (Figures 6.1, 6.2 and 6.3), and in the historical record they are most commonly tide-gauge measurements (Figure 6.4) or high-water marks on man-made coastal structures. A strandline, or a sequence of strandlines, is a relative sea-level record that potentially represents both real and apparent sea-level changes (Figure 6.5): (6.1) Real sea-level changes are absolute vertical movements of the ocean’s surface and may be local to worldwide in extent; if worldwide, they are called eustatic changes. FIGURE 6.5 Relative, apparent, and real sea-level changes. A, Late Pleistocene. B, Holocene. All sea-level records (strandlines or tide-gauge measurements) are relative, which means they potentially represent both apparent and real sea-level changes (relative=real +apparent). Apparent sea-level changes are the inverse of the vertical crustal (or ground) movements that produce them and, therefore, are the focus of all coastal tectonic studies. The apparent sea-level history is obtained by algebraically or graphically subtracting the real sea-level history from the relative record of marine strandlines. In effect, the real sea-level history is a composite tectonic datum. The real sea-level history is obtained by subtracting apparent sea level changes from a relative sea-level record. In this case, the uplifting coastline is a moving sea-level datum. E is the present elevation of a strandline; e is the original elevation of a strandline; D is the vertical displacement (D=E−e); A is the age of a strandline; R is the crustal displacement rate (R=D/A).
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Active Tectonics: Studies in Geophysics FIGURE 6.6 Pleistocene sea-level fluctuations and origin of emergent Pleistocene strandlines. Emergent strandlines simultaneously record tectonic uplift and major sea-level highstands. The rising coastline is a moving strip chart on which sea-level highstands are recorded sequentially as strandlines whose ages increase with elevation. The slope (R) of the diagonal line connecting each highstand to the elevation of its strandline is the average uplift rate. If the uplift rate was constant, the uplift lines for all strandlines are parallel. Strandlines formed during lowstands are usually destroyed by subsequent sea-level fluctuations and rarely appear in the emergent geologic record. Strandlines younger than 60 ka appear above sea level only where the uplift rate is greater than 1 m/ka. The sea-level fluctuation curve was derived from a sequence of U-series dated coral-reef strandlines on the Huon Peninsula, Papua New Guinea (Figure 6.2B) by subtracting tectonic uplift from the relative strandline record (Figure 6.5A). Sea-level curve modified from Chappell (1983); oxygen-isotope stages (1–9) from deep-sea cores (Shackleton and Opdyke, 1973). Quaternary sea-level history was characterized by periodic eustatic fluctuations of 100–150 m caused by the advance and retreat of continental glaciers. Apparent sea-level changes are not real but result from and are the inverse of vertical ground movements. Consequently, apparent sea-level changes are only local or regional in extent. The primary task in coastal tectonics is separating the apparent component from the real component of a relative sea-level record. This is done graphically or algebraically by subtracting the real sea-level history from a relative sea-level record: (6.2) In effect, the real sea-level history is a fluctuating tectonic datum to which each strandline (tectonic marker) must be correlated by age (Figures 6.5, 6.6, and 6.7). The real sea-level history is obtained by subtracting apparent sea-level changes (the inverse of vertical crustal movements) from a detailed and well-dated relative sea-level record (Figure 6.5): (6.3) In this case, the tectonic history is an absolute sea-level datum and can be either moving or stationary. A rapidly and steadily rising coastline is the best datum for measuring major, long-term sea-level fluctuations (Figures 6.5A and 6.6) (Chappell, 1983), whereas a stable coastline is the best datum for measuring minor, short-term sea-level changes (Figures 6.5B and 6.7) (Bloom, 1970; Scholl et al., 1970). Generally, the greatest uncertainties in most long-term sea-level histories derived from strandline data stem from the simplifying assumption of constant uplift. However, on a tilted coastline where Pleistocene strandlines converge, constant long-term uplift at a particular locality can be demonstrated empirically (Chappell, 1983). The greatest uncertainties in short-term sea-level histories stem from the assumption of coastal stability. Where short-term stability at a coastal locality cannot be demonstrated by independent geodetic data, subtle sea-level changes can be expressed in only relative terms. An additional complication is FIGURE 6.7 Holocene sea-level changes and origin of emergent Holocene strandlines. In contrast to Pleistocene strandlines, emergent Holocene strandlines represent discrete uplift events or storm events, not sea-level fluctuations. The highest strandline correlates with the tangent point at the sea-level inflection between 7 ka and 5 ka BP. All lower strandlines represent two paleoshorelines, one occupied during the transgressive phase and one during the regressive phase (see Figure 6.5B for relative sea-level changes). The sea-level curve is general and may vary slightly from region to region owing to minor geoidal distortions.
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Active Tectonics: Studies in Geophysics that different measurement techniques (e.g., geodetic, tide gauge) occasionally yield conflicting results for short-term sea-level changes (Brown, 1978). Pleistocene Strandlines Along steep coastal slopes in uplifting areas Pleistocene strandlines occur most commonly as narrow (<1 km), steplike terraces within a few hundred meters above present sea level. Consequently, a vertical sequence of uplifted strandline terraces resembles a flight of stairs (Figures 6.1 and 6.2). Each strandline terrace consists of a virtually horizontal erosional or depositional platform backed by a steep sea cliff along its landward margin. The shoreline angle, the intersection of the relict platform and sea cliff, closely approximates the location and elevation of the abandoned marine shoreline and is the linear structural marker depicted on most longitudinal profiles of deformed strandlines (for examples see Figures 6.8A, 6.13, 6.17, 6.20, and 6.21). The shoreline platform, also referred to as the terrace platform, is the planar tectonic marker depicted on most cross-sectional profiles and detailed isobase maps of deformed strandlines (Figures 6.2 and 6.3A; also see Figures 6.19 and 6.21 below). Commonly, a thin (1–3 m) veneer of shallow-water sand, gravel, and cobbles, which locally contains fossil marine shells used for dating, overlies the terrace platform. Where subaerial slope degradation is rapid, a seaward-thinning prism of alluvium derived from upland streams and abandoned sea cliffs overlies the platform and its associated marine sediments and buries the shoreline angle (Figure 6.2A). In these areas the geomorphic expression of successively higher (older) Pleistocene strandline terraces is progressively subdued, and the position and configuration of both the shoreline angle and the terrace platform must be derived from borehole or shallow seismic data (see Figure 6.21 below) (Bradley and Griggs, 1976; Lajoie et al., 1979b). However, in some areas the progressively greater degradation of successively older relict sea cliffs provides a means of dating emergent strandlines (Figure 6.2A) (Hanks et al., 1984). Along gentle coastal slopes in slowly uplifting areas emergent strandlines commonly consist of broad (1–10 km) terrace platforms backed by low sea cliffs obscured by relict beach ridges and dune fields (Hoyt and Hails, 1974; Lajoie et al., 1979a). A general consensus has developed over the past two decades that a flight of emergent Pleistocene strandlines is the geologic record of periodic glacio-eustatic sea-level highstands superimposed on a rising coastline (Figure 6.6) (Broecker et al., 1968; Mesolella et al., 1969; Matthews, 1973). In this model, a rising coastal land-mass is a moving strip chart on which brief sea-level highstands were successively recorded as depositional or erosional strandlines. Strandlines also formed during sea-level lowstands (Emery, 1958; Lewis, 1971a,b, 1974; Ridlon, 1972), but along uplifting coastlines most of these strandlines were destroyed by wave erosion during subsequent sea-level fluctuations and, therefore, rarely appear in the emergent marine record (Figure 6.6). An important exception is found in the deeply incised sequence of emergent strandline terraces on the Huon Peninsula of Papua New Guinea, where sea-level lowstands are recorded as deltaic accumulations preserved beneath a protective cover of coral reefs deposited during subsequent sea-level highstands (Chappell, 1974a, 1983). Along many subsiding coastlines, strandlines formed during sea-level lowstands probably dominate the submergent geomorphic record but are difficult to distinguish from strandlines formed during highstands (Moore and Fornari, 1984). The most detailed tectonic datum for deriving uplift from emergent Pleistocene strandlines on coastlines throughout the world is the paleosea-level curve obtained by subtracting well-documented constant tectonic uplift from the relative sea-level record of emergent coral-reef strandlines on the Huon Peninsula of Papua New Guinea (Figures 6.2B and 6.6) (Veeh and Chappell, 1970; Bloom et al., 1974; Chappell, 1983). There, uranium-series dates on fossil corals yield a paleosea-level curve back to about 340 ka before present (BP) that agrees with longer, but less detailed, sea-level curves derived from emergent coral-reef strandlines on the island of Barbados in the West Indies (Broecker et al., 1968; Mesolella et al., 1969; Matthews, 1973; Bloom et al., 1974; Bender et al., 1979) and from oxygen-isotope data from deep-sea cores (Shackleton and Opdyke, 1973). The uplift rate (maximum 4.0 m/ka) of the Huon Peninsula was derived from the elevation of the prominent 120-ka strandline that formed 2–10 m above present sea level during the last major interglacial sea-level highstand (Thurber et al., 1965; Veeh, 1966; Ku et al., 1974; Neumann and Moore, 1975; Marshall and Thom, 1976; Stearns, 1976; Harmon et al., 1978; Shubert and Szabo, 1978; Szabo et al., 1978; Szabo, 1979). The most important features of the New Guinea paleosea-level curve as a tectonic datum are the periodic interglacial highstands at approximately 100 ka intervals and the successively lower interstadial highstands at approximately 20 ka intervals over the past 120 ka. Virtually all Pleistocene strandlines along emergent coastlines throughout the world were formed during these brief paleosea-level highstands. The longer paleosea-level records on Barbados and in the deep-sea cores indicate that the main 100-ka cycle of interglacial high-
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Active Tectonics: Studies in Geophysics stands extends back to at least 700 ka (Shackleton and Opdyke, 1973; Bender et al., 1979). Morner (1976, 1983) argued that no locally derived Pleistocene sea-level curve should be used as a worldwide tectonic datum because the sea surface, an approximate geoid or equipotential gravitational surface, is grossly distorted by gravitational fluctuations on a time scale of 10–100 ka. If correct, this supposition implies that major, glacially induced changes in ocean volume were expressed by neither synchronous nor uniform sea-level fluctuations. However, similar absolute ages of Pleistocene strandlines on emergent coastlines throughout the world and, more importantly, similar elevations of sea-level highstands on several independently derived late Pleistocene sea-level curves demonstrate that, within reasonable margins of uncertainty (±5 ka and ±10 m; Stearns, 1976, 1984; Harmon et al., 1979; Chappell, 1983), major sea-level fluctuations were synchronous and relatively uniform over at least the past 300 ka, and most likely much longer (Veeh and Valentine, 1967; Bloom et al., 1974; Konishi et al., 1974; Ku and Kern, 1974; Moore and Samayajulu, 1974; Chappell and Veeh, 1978; Harmon et al., 1978; Marshall and Launay, 1978; Bender et al., 1979; Dodge et al., 1983; Ward, 1985; also see references on elevation of 120-ka highstand in preceding paragraph). Also, reasonable graphical correlations of undated strandline sequences in Japan (see Figure 6.14 below) (Miyoshi, 1983), New Zealand (W.Bull, University of Arizona, personal communication, 1983), and California (Hanks et al., 1984) with 120-ka to 40-ka highstands on the New Guinea paleosea-level curve indicate that strandlines from all four areas were synchronous and had a common datum. An apparent exception is found along the southeast Atlantic coast of North America, where the ages, but not the relative heights, of late Pleistocene highstands agree with paleosea-level data from New Guinea and other parts of the world (Belknap, 1979; Cronin et al., 1981). However, the discrepancies of 20–30 m found along this passive continental margin may represent subtle tectonic or isostatic crustal movements not regional sea-level variations. In any event, although rapid geoidal distortion is a potential problem that requires further investigation, most existing data indicate that geoidal distortion could not have been significant over late Pleistocene time. Consequently, there is no compelling reason not to use the New Guinea sea-level curve worldwide as a first-order tectonic datum, at least over the past 120 ka. Because virtually all emergent Pleistocene strandlines were formed during sea-level highstands (Figure 6.6), the task of dating a particular strandline is reduced to correlating it with a specific peak on the New Guinea paleosea-level curve using one or more absolute (isotopic) or relative (geomorphic, paleontologic, and chemical) dating techniques (Chappell and Veeh, 1978; Kennedy et al., 1982; Sutherland, 1983; Hanks et al., 1984). Uranium-series analyses of fossil corals yield the most reliable absolute dates for Pleistocene strandlines (Broecker and Bender, 1973; Harmon et al., 1979), but because fossil corals are virtually restricted to tropical latitudes, less-reliable dating techniques must be used at higher latitudes. Frequently, only one or two Pleistocene strandlines in an emergent sequence can be dated independently, but the ages of other strandlines can be inferred from the resultant uplift rate (Chappell and Veeh, 1978; Hanks et al., 1984; Ward, 1985). Occasionally, an entire sequence of undated late Pleistocene strandlines can be correlated with the paleosea-level curve by trial and error if different, but constant, uplift rates are assumed; the best graphical correlation of strandline elevations and paleosea-level highstands simultaneously yields the most reasonable uplift rate and strandline ages (see Figure 6.6). In effect, the uplift rate itself is a useful dating tool, even if not independently derived. However, in some complexly deformed coastal areas no strandlines can be dated confidently using this graphical technique (Weber, 1983), which suggests that the uplift rate varied with time. Generally, for strandlines dated between 700 ka and 200 ka BP by either absolute or relative techniques the uncertainty in correlation with the Pleistocene paleosea-level curve is at least one major glacioeustatic cycle (±100 ka), and for strandline dated between 120 ka and 30 ka the uncertainty is commonly one minor cycle (±20 ka) (Harmon et al., 1979; Stearns, 1984). The best-preserved and most reliably dated Pleistocene strandlines along most emergent coastlines correlate with the 120-ka, 104-ka, and 82-ka sea-level highstands (see Figure 6.11B below). These three highstands correlate with deep-sea oxygen-isotope substages 5e, 5c, and 5a, respectively (Shackleton and Opdyke, 1973; Chappell, 1983), and their strandlines are often referred to by these alphanumeric designations. Strandlines that correlate with younger (lower) highstands appear in the emergent geologic record only where uplift rates are sufficiently high (>0.3–1.0 m/ka) to elevate them above present sea level (Figure 6.6). Along most erosional coastlines some emergent Pleistocene strandlines are missing or are laterally discontinuous owing to sea-cliff retreat during the formation of younger strandlines (see Figures 6.17B and 6.20A below). Generally, strandlines older than 400 ka are poorly preserved owing to subaerial slope degradation and stream incision; exceptions are found on stable or slowly rising coastlines where strandlines as old as 2.4 m.y. are commonly preserved
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Active Tectonics: Studies in Geophysics (Ward, 1985). In areas of rapid uplift (greater than 4 m/ka), subaerial erosional processes are greatly accelerated and strandlines older than 120 ka are rarely preserved. However, in areas of rapid uplift the fine structure of late Pleistocene sea-level history is most clearly and completely recorded by strandlines younger than 120 ka (Chappell, 1983). Holocene Strandlines Radiocarbon dates on fossil wood, peat, and shell from sedimentary deposits on continental shelves and in shallow coastal embayments throughout the world indicate that sea level stood about 100–150 m below its present position during the last glacial maximum between 20 ka and 15 ka BP, and then rose rapidly to within 4–6 m of its present position by about 7–5 ka BP as the major ice sheets retreated at the end of Pleistocene time (Bloom, 1977). Although sea-level changes over the past 5 ka may have varied regionally owing to minor geoidal distortion and hydro-isostatic crustal loading (Clark et al., 1978), the most reliable data indicate that sea level has not fluctuated significantly (±2 m), nor has it been higher than its present position, over this period of time (Figure 6.7) (Thom et al., 1969; Bloom, 1970; Scholl et al., 1970; Omoto, 1979; Faure et al., 1980; Thom and Roy, 1983; Sneh and Klein, 1984; Gibb, in press). Consequently, Holocene strandlines, unlike Pleistocene strandlines, do not represent eustatic sea-level fluctuations but rather episodic crustal movements or major storm events. However, the highest strandlines in most Holocene strandline sequences usually represent the sudden and drastic decrease in postglacial sea-level rise between 7 ka and 5 ka BP (Figure 6.7). Along erosional coastlines emergent Holocene strandlines form and survive only where crustal uplift is sufficiently rapid (>1–2 m/ka) to offset the destructive effects of subsequent wave erosion. However, along prograding coastlines depositional strandlines survive even where there is little or no crustal uplift (Figure 6.3B) (Schofield, 1973, Garrick, 1979). Most emergent Holocene strandlines of both erosional and depositional origin are similar in configuration to their Pleistocene counterparts but usually are smaller in scale, finer in geomorphic expression, and better preserved (Figure 6.3). Also, because of their relatively youthful age, most Holocene strandlines lie within tens rather than hundreds of meters above present sea level, even along the most rapidly uplifting coastlines. In protected embayments and near the mouths of large streams and rivers, Holocene strandlines commonly consist of bouldery to sandy beach ridges that are relatively imprecise paleosea-level indicators (Figure 6.3B) (Schofield, 1973; Garrick, 1979). However, along many wave-resistant coastlines with small tidal ranges (less than 1 m) emergent Holocene strandlines consist of horizontal solution notches and faint waterlines on rocks and sea cliffs that are extremely precise paleosea-level indicators (Machida et al., 1976; Pirazzoli et al., 1982). Along some recently uplifted coastlines, strandlines consist of horizontal bands of fossil intertidal sessile organisms such as barnacles, which are also precise paleosea-level indicators (Sawamura, 1953) even in areas with large tidal ranges (Plafker, 1965). In the historical record, emergent Holocene strandlines most commonly consist of points on tide-gauge records (Figures 6.4 and 6.9B) (Hicks and Crosby, 1974; Berrino et al., in press), but in a few areas they consist of waterlines and burrows of intertidal organisms on recently uplifted or down-dropped man-made coastal structures (see Figure 6.9A below) (Grant, 1970; Flemming, 1972; Berrino et al., in press). Along many rapidly uplifting coastlines, prehistoric Holocene strandlines are similar to strandlines produced by abrupt vertical crustal movements associated with major historical earthquakes (Sugimura and Naruse, 1954; Plafker, 1965; Wellman, 1969; Matsuda et al., 1978). This similarity suggests that all emergent Holocene strandlines are of coseismic origin. However, along stable or steadily uplifting coastlines, both erosional and depositional strandlines probably form during infrequent major storms (Schofield, 1973; Garrick 1979; Hillaire-Marcel, 1980). A major uncertainty in deriving a history of paleoseismicity from a sequence of emergent Holocene strandlines is that coseismic and storm-produced strandlines may be indistinguishable. Consequently, if coseismic and storm-produced strandlines occur together (Sandweiss and Rollins, 1981; Lajoie et al., 1982a), the number of past earthquakes could be overestimated from the geologic record. On the other hand, wave erosion may remove coseismic strandlines from the geologic record (see Figure 6.27B below) (Plafker et al., 1981), in which case the number of earthquakes represented by a sequence of emergent Holocene strandlines could be underestimated. Because of minor regional or subregional geoidal distortions, local hydro-isostatic adjustments, variable oceanographic conditions (currents, water temperature, salinity), and climatic fluctuations, there is probably no locally derived late Holocene sea-level curve that can be used as a precise universal tectonic datum for measuring minor (less than 2 m) vertical crustal movements (Clark et al., 1978; Newman et al., 1978). At best, a locally derived curve is a regional tectonic datum. However, data from coastlines throughout the world indicate that relative sea-level changes greater than 2–4 m over the past 6 ka are either apparent (caused by tectonic and isostatic crustal movements or sediment
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Active Tectonics: Studies in Geophysics compaction) or spurious (caused by reworking or contamination of dated fossil materials), not real. Therefore, present sea-level or a generalized sea-level curve (Figure 6.7) is often used as an approximate tectonic datum for measuring late Holocene crustal movements, especially where rates of deformation are very high (see Figures 6.25, 6.26, and 6.27 below) (Matsuda et al., 1978; Plafker and Rubin, 1978; Plafker et al., 1981). Most Holocene strandlines in the geologic record are dated directly or indirectly by radiocarbon techniques. However, the sudden and drastic decrease in the sea-level rise between 7 ka and 5 ka BP occasionally leads to incorrect or misleading radiocarbon dates on fossil materials (particularly marine shells) from emergent Holocene strandlines. This potential problem exists because all but the highest strandline in a sequence of emergent Holocene strandlines represent two different paleoshorelines—a transgressive shoreline older than about 6 ka and a regressive shoreline younger than 6 ka (Figure 6.7). Consequently, marine shells deposited on the older, transgressive shoreline could be reworked into sediments deposited on the younger, regressive shoreline. On the other hand, the two-part configuration of the Holocene sea-level curve provides an independent means of approximating the ages of emergent Holocene strandlines and deriving minimum uplift rates. On a plot of strandline elevation versus sea-level history (Figure 6.7), the straight line connecting the elevation of the highest strandline on the vertical axis and the tangent point at the break on the sea-level curve represents the average uplift since that strandline formed. The coordinates of the tangent point are the tentative age and original elevation of the highest strandline, and the slope of the line is the tentative average uplift rate. If constant uplift is assumed, interpolation of the uplift rate yields age estimates for any lower strandlines (see Figure 6.28 below) (Wellman, 1969). Because the highest Holocene strandline on most coastlines represents the break in sea-level rise, its age is usually between 7 ka and 5 ka, depending on the uplift rate; because the break is rounded, the age of the highest strandline increases from 5 ka to 7 ka as the uplift rate increases. However, if subsequent cliff erosion destroyed the 7–5-ka strandline, the highest surviving strandline would be somewhat younger. Obviously, other dating techniques are needed to test for this possibility. STRANDLINE DISPLACEMENT AND DEFORMATION Regional patterns of vertical strandline displacement (uplift or subsidence) and deformation (folding and faulting) commonly reflect primary tectonic processes (subduction and rifting) related directly to horizontal plate motions. Important exceptions are found in the northernmost continental areas of North America and Europe where regional patterns of latest Pleistocene and Holocene strandline displacement reflect major but transitory isostatic responses of the crust and mantle to geologically recent glacial unloading. Local patterns of strandline displacement and deformation generally involve minor tectonic structures (folds, faults, horsts, and grabens) that reflect secondary tectonic processes related to subregional stress patterns. Available data indicate that most secondary and some primary tectonic deformation recorded by marine strandlines extends no farther than a few tens of kilometers inland from the coastline. In many areas, vertical ground displacements on a local scale also reflect nontectonic processes such as volcanic tumescence (Kaizuka et al., 1983; Berrino et al., in press), isostatic adjustments due to sediment accumulation (Fisk and McFarlan, 1955), and sediment compaction due to natural processes (Atwater et al., 1977) or fluid extraction (Poland, 1971; Buchanan-Banks et al., 1975). In many cases, especially those involving minor vertical displacements, it is difficult, if not impossible, to distinguish tectonic from nontectonic apparent sea-level changes. It is noteworthy that on both local and regional scales sequences of marine (and lacustrine) strandlines constitute the longest and most detailed and areally extensive records of late Quaternary crustal deformation. Indeed, the magnitudes and rates of tectonic, isostatic, and volcanic processes derived from deformed and displaced strandlines are commonly used as references for comparing and interpreting tectonic data from less complete or shorter records in other geological environments. Vertical Displacement Because marine strandlines define horizontal lines and planes (where curved or closed), they record vertical crustal displacements (uplift or subsidence) more clearly than horizontal displacements. Net vertical displacement (D) is the difference between the present elevation (E) and the original elevation (e) of a strandline and is also the product of the strandline’s age (A) and the average displacement rate (R) (Figure 6.5): (6.4) The average displacement rate (R) is the displacement (D; E−e) of a strandline divided by its age (A): (6.5) An isobase is a locus of points on a planar tectonic marker (a marine platform or a plane defined by a curved or closed strandline) along which the vertical
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Active Tectonics: Studies in Geophysics displacement rate (R) is equal (see Figures 6.12, 6.15A, 6.16A, 6.18, and 6.21 below). The history of vertical crustal displacements at a coastal locality is derived graphically by plotting the displacement (D) of each independently dated strandline as a function of its age (A) (Figure 6.5). The resultant locus of points is the apparent sea-level history, which, as stated previously, is the inverse of the displacement history. If this locus of points defines a straight line, the displacement rate (R), which is the inverse of the slope of the line, was constant. If it does not define a straight line, the displacement rate was variable (see Figures 6.9A and 6.10B below). Commonly, only one strandline in a sequence can be dated independently, and, therefore, only an average displacement rate can be derived from a relative sea-level record. A plot of strandline elevation (E) as a function of age (A) yields the relative sea-level history, which is the inverse of the approximate displacement history where vertical displacements are large compared to sea-level changes. Frequently, uplift histories based on highly emergent Holocene strandlines (see Figures 6.25, 6.26, and 6.27 below) or very old Pleistocene strandlines (Bender et al., 1979) are approximated by relative sea-level curves. In other words, present sea level can be used as an approximate datum. The presence of emergent marine strandlines along most active-tectonic coastlines suggests that crustal uplift is more common than subsidence. However, subsidence is probably underestimated merely because most geologic evidence for downward crustal movements is buried in sedimentary basins (Atwater et al., 1977) or is submerged offshore (Lewis, 1974). Many studies of long-term crustal movements in coastal areas focus on uplift mainly because the emergent strandline record is better exposed and easier to interpret than the submergent record. Most long-term crustal movements recorded by marine strandlines reflect sustained tectonic deformation along active plate boundaries, but the most rapid known rates of vertical crustal displacement reflect transitory volcanic tumescence and glacio-isostatic rebound. Both processes are noteworthy primarily for comparative purposes, but also because they are, themselves, related to tectonic processes in various ways. Volcanic Displacements The highest known rates of sustained vertical ground displacement exceed 100 mm/ yr and are recorded by marine strandlines on the flanks of active insular and coastal volcanoes. For example, on the island of Iwo Jima, the tip of a large volcano in the western Pacific Ocean, radiometrically and historically dated emergent strandlines yield an average uplift rate FIGURE 6.8 Recent volcanic uplift of Iwo Jima, a volcanic island 1200 km south of Tokyo, Japan. A, Longitudinal strandline profiles. The strandline labeled 1779 was the active shoreline mapped in 1779 by the crew of the English explorer, Captain Cook. The age and maximum elevation of this strandline yield an average uplift rate of 200 mm/yr, which is similar to the rate of 170–240 mm/yr derived from radiocarbon dates of 0.5–0.7 ka on the 110-m strandline. The well-defined strandlines were formed by wave action during brief pauses in uplift or during periodic storms. B, Average uplift rates derived from 110-m (0.5–0.7 ka) and 40-m (0.2 ka) strandlines. Tide-gauge data record variable uplift that reached 800 mm/yr over the past 80 yr. Both modified from Kaizuka et al. (1983). of 200 mm/yr over the past 800 yr (Figure 6.8) (Kaizuka et al., 1983). However, tide-gauge data suggest that uplift rates on the island probably fluctuated between 100 and 800 mm/yr over this period of time (Figure 6.8B), during which only minor phreatic eruptions are known to have occurred (Corwin and Foster, 1959). An even longer record of vertical ground displacements related to volcanic processes is found in the Phlegraean Fields caldera on the Mediterranean coast near Naples, Italy, where historically dated strandlines on man-made structures document alternating subsidence and uplift that averaged 12 mm/yr over the past 2 ka (Figure 6.9) (Berrino et al., in press). However, uplift and subsidence exceeded 150 mm/yr for a few decades before and after a minor eruption in A.D. 1538. Tide-
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Active Tectonics: Studies in Geophysics gauge records at Pozzuoli, a coastal suburb of Naples, yield uplift rates of 350 and 500 mm/yr for two brief periods between 1970 and 1983 (Figure 6.9B), which suggests that another eruption is imminent. However, the data from Iwo Jima indicate that rapid uplift and even short surges of extremely rapid uplift are not by themselves definite indicators of imminent eruption. Glacio-Isostatic Displacements Highly emergent latest Pleistocene and Holocene marine (and lacustrine) strandlines in the deglaciated areas of northern North America and Europe record the highest known rates of regional vertical crustal displacement (Figure 6.10). Around Hudson Bay in Canada (Farrand and Gajda, FIGURE 6.9 Recent volcanic displacements at Pozzuoli, a coastal suburb of Naples, Italy, in the Phlegraean Fields caldera. A, Vertical displacement based on historical data such as water marks on Roman buildings. Mount Nuovo, a small cinder cone nearby, erupted in A.D. 1538, about 40 yr after the uplift rate increased from 10 mm/yr to about 100 mm/yr. The general subsidence that followed this small eruption has now reversed. B, Tide-gauge data at Pozzuoli Harbor from 1970 to 1983. During two brief periods (1970–1973 and 1982–1983) uplift reached 800 mm/ yr. Because of this rapid uplift, increased fumerole activity, and ground cracking, parts of Pozzuoli have been evacuated in anticipation of another eruption. Modified from Berrino et al. (in press). FIGURE 6.10 Postglacial isostatic rebound curves derived from dated marine and lacustrine strandlines. A, Relative sea-level changes (minimum apparent sea-level changes), Gulf of Bothnia, Scandinavia. Modified from Lundqvist (1965). B, Apparent sea-level changes, Hudson Bay, North America. Modified from Farrand and Gajda (1962). In both areas crustal movement is the inverse of the curves. Regional crustal uplift exceeded 100 m/ka shortly after deglaciation in both areas. 1962; Andrews, 1970; Peltier and Andrews, 1983) and the Gulf of Bothnia in Fennoscandia (Lundqvist, 1965; Donner, 1965, 1980), which were the regions of thickest ice accumulation during the last major glacial advance, rates of postglacial isostatic rebound reached 100 m/ka (100 mm/yr) between 13 ka and 7 ka BP and decreased quasi-exponentially to the present rates of 6–9 mm/yr (Barnett, 1966; Balling, 1980). Even these rates of residual glacio-isostatic rebound are extremely high compared with most rates of vertical tectonic displacement. Along the margins of these deglaciated areas, Holocene strandlines and tide-gauge data record complex histories of rapid subsidence as well as uplift, which probably reflect the collapse and inward migration of a crustal forebulge peripheral to the retreating ice front (Grant, 1980; Morner, 1980; Clague et al., 1982).
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Active Tectonics: Studies in Geophysics FIGURE 6.21 Closely related folding and faulting recorded by deformed 82-ka strandline at Half Moon Bay, California. Structural contours on wave-cut platform depict drag folds that plunge obliquely into and away from the vertical fault plane. Orientation of folds indicates predominant right-lateral fault movement, which is consistent with regional fault movements. Longitudinal profiles of strandline and faulted platform record gentle warps and vertical offset across fault. Shoreline angle of strandline does not intersect fault plane so local slip vector and slip rate are not known. Vertical separation of wave-cut platform across fault revealed by platform profiles on fault plane (dotted lines; E, east; W, west). Structural contours derived mainly from shallow borehole and seismic refraction data. Modified from Lajoie et al. (1979b). island (Figure 6.15) (Tamura, 1979). If this fault is vertical, the 120-ka strandline is offset about 75 m, which yields a long-term slip rate of 0.6 m/ka. The 6-ka strandline is offset about 4 m, which yields a similar short-term slip rate of 0.7 m/ka. Another excellent example of continuous fault slip recorded by progressively greater offset of successively older marine strandlines is found near the town of Ventura on the southern California coastline (Figure 6.19B) (Sarna-Wojcicki et al., 1986). There, the wave-cut platform of the 40-ka strandline (Kennedy et al., 1982) is vertically offset about 45 m across a high-angle reverse fault, which yields an average slip rate of 1.1 m/ka. Nearby, a stream-cut platform graded to an emergent 3.5-ka strandline is offset 4 m across the same fault, which yields a similar shorter-term slip rate of 1.2 m/ka. If the size of an average coseismic slip event is assumed or is known from a historical seismic event, it can be divided into the slip rate to yield the average earthquake recurrence interval. However, where discrete slip events are recorded geologically, earthquake recurrence can be determined more directly. For example, lenses of scarp-derived Holocene talus preserved beneath the upthrown block of the high-angle reverse fault near Ventura, California (Figure 6.19B), record at least five discrete slip events during the past 3.5 ka (Figure 6.22A) (Sarna-Wojcicki et al., 1986). If these slip events were of coseismic origin, the average vertical displacement was FIGURE 6.22 Coseismic slip events on faults recorded by scarp-derived talus. A, Stream-cut platform graded to 3.5-ka marine strandline offset 4 m across Javon Canyon fault near Ventura, California (see Figure 6.19B). At least five discrete slip events (0.5–1.4 m) are recorded by lenses of scarp-derived talus preserved below the upthrown block. If these slip events are coseismic, the average earthquake recurrence interval for this fault is about 700 yr. Modified from Sarna-Wojcicki et al. (1986). B, Vertical component of fault displacement recorded by offset of 104-ka wave-cut platform at Point Año Nuevo, California. At least nine discrete slip events are recorded by lenses of scarp-derived talus and faulted sediment preserved on the wave-cut platform beneath the upthrown block. If slip events are coseismic, the average earthquake recurrence interval is about 12 ka. It should be stressed, however, that average recurrence intervals have little meaning if total fault movement occurred during a short period of time (dotted line). Independently derived ages for each event are needed to resolve this potentially serious problem. Modified from Weber and Cotton (1981). 0.8 m and the average earthquake recurrence interval was 700 yr. An even longer history of discrete slip events is recorded where a 104-ka marine platform is offset across a strand of a complex fault zone at Point Año Nuevo about 100 km south of San Francisco on the central California coastline (Lajoie et al., 1979b). This wave-cut platform
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Active Tectonics: Studies in Geophysics is offset 5.5 m (dip-slip component), which yields an average long-term slip rate of 0.05 m/ka. More importantly, however, scarp-derived talus that accumulated on the marine platform below the periodically rejuvinated fault scarp records at least nine slip events over the past 104 ka (Figure 6.22B) (Weber and Cotton, 1981). If coseismic, these slip events, which averaged 0.6 m, yield an average earthquake recurrence interval of about 12 ka. It should be stressed, however, that if total fault movement was restricted to relatively short periods of time (Figure 6.22B), average recurrence intervals are of little use in predicting the time of the next event. Each event must be dated independently to make meaningful predictions of future events. Lateral Fault Movement As indicated previously, lateral fault movement is rarely recorded by marine strandlines, even where strandlines clearly cross active strike-slip faults. Usually, the amount of offset is ambiguous because of irregularities in or poor preservation of the paleoshoreline. For example, at Point Año Nuevo on the central California coastline, two Pleistocene strandlines that cross a major right-lateral fault system appear to be offset across several fault strands (Weber and Lajoie, 1977; Weber and Cotton, 1981). However, because these strandlines must be projected considerable distances (up to 1.0 km) at low angles across some of the fault strands, the individual and cumulative offsets of each strandline are too uncertain to yield meaningful rates of lateral fault movement. In some areas, the sense but not the amount of lateral fault movement is expressed by the orientation of drag folds recorded by marine strandlines. For example, the 82-ka strandline at Half Moon Bay about 60 km northwest of Point Año Nuevo on the central California coastline records broad crustal warps on both sides of a major fault strand within the coastal fault system (Figure 6.21) (Lajoie et al., 1979b; Kennedy et al., 1981). Structural contours on the warped wave-cut platform define a broad syncline that plunges obliquely into the vertical fault plane from the east and a broad anticline that plunges away from the fault plane to the west. These structural relationships suggest that the syncline and anticline are drag folds related to right-lateral fault movement, which is consistent with regional fault movements. Unfortunately, the strandline of this deformed marine platform does not intersect the fault plane and, therefore, neither the local slip vector nor the slip rate is known. However, where latest Pleistocene stream courses cross the warped marine terrace northeast of the fault they are deflected toward the axis of the syncline (Figure 6.21), which suggests that fault movement and related crustal warping were continual over the past 82 ka. COSEISMIC UPLIFT AND EARTHQUAKE RECURRENCE In several seismically active coastal areas historical coseismic uplift has produced conspicuous emergent marine strandlines 1–15 m above sea level. The best documented examples of historical coseismic strandlines are found in Japan (Sugimura and Naruse, 1954; Nakamura et al., 1965; Ota et al., 1976; Matsuda et al., 1978; Shimazaki and Nakata, 1980), New Zealand (Wellman, 1969; Stevens, 1973), Alaska (Plafker, 1965, 1972), Chile (Plafker and Rubin, 1967; Plafker and Savage, 1970; Plafker, 1972), and Iran (Page et al., 1979). The historical coseismic strandlines in these areas suggest that similar higher (older) Holocene strandlines in these and other active-tectonic areas are also of earthquake origin. If this interpretation is correct, a sequence of emergent Holocene strandlines is a physical record of past earthquakes from which it should be possible to predict the size and date of the next seismic event. If the average uplift rate is constant, coseismic uplift events may follow a time-predictable or displacement-predictable pattern (Figure 6.23) (Shimazaki and Nakata, 1980). In a time-predictable pattern the time between events is proportional to the size of the preceding event, and, therefore, the date, but not the size, of the next event can be predicted. In a displacement-predictable pattern the time between events is proportional to the size of the succeeding event, and the size of the next event can be predicted for any future date, but that date cannot be predicted. Of course, if the period and size of uplift events are regular, both can be predicted for the next event. In practice, however, strandline elevations and dates are usually too variable or imprecise to fit any pattern exactly, and, therefore, only average displacements and recurrence intervals can be derived from sequences of emergent Holocene strandlines. In some cases, the uplift rate is so variable that no reasonable predictions can be made. A few examples of coseismic uplift from Japan, Alaska, and New Zealand illustrate some of the possibilities and limitations in deriving detailed seismic histories from sequences of emergent Holocene strandlines. Japan The Japanese archipelago, which lies along the leading, overthrust margin of the Eurasian tectonic plate in the western Pacific Ocean, is one of the most seismically
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Active Tectonics: Studies in Geophysics FIGURE 6.23 Theoretical patterns of coseismic uplift (or fault movement). A, If average uplift rate is constant, the incremental coseismic uplift events may follow either a time-predictable or a displacement-predictable pattern. In a time-predictable pattern the time between events (T) is proportional to the size of the preceding displacement (PD). In this case, the time, but not the size, of the next event can be predicted. In a displacement-predictable pattern the time between events is proportional to the size of the succeeding displacement (SD). In this case, the size of the next event can be predicted for any future date, which cannot be predicted. In practice, strandline elevations and ages are usually too uncertain to reveal precise patterns of uplift. Modified from Shimazaki and Nakata (1980). B, If a displaced marker (e.g., wave-cut platform) is dated but individual events are not dated independently, their intervals and age ranges can be estimated from the average displacement rate. However, if total displacement is highly irregular (dotted line), average displacement intervals, rates, and age values are meaningless. Independent ages of events are needed to establish displacement pattern. active areas on Earth (Tarr, 1974). Great intraplate megathrust earthquakes occur along the major subduction zone that forms the eastern boundary of the archipelago, and smaller, but still potentially destructive, interplate earthquakes occur along its complexly faulted eastern boundary. Particularly good examples of coseismic strandlines are found on the Boso Peninsula, the Muroto Peninsula, and Kikai Island along the overthrust Pacific Coast and on Awashima Island and the Ogi Peninsula of Sado Island along the block-faulted Japan Sea Coast (Figure 6.24). Four emergent Holocene strandlines (Numa I-IV) form prominent terraces around the southern tip of the Boso Peninsula about 80 km south of Tokyo (Figures 6.3A, 6.24, and 6.25) (Sugimura and Naruse, 1954; Matsuda et al., 1978). Isobases on the planes defined by these curved strandlines reveal progressive uplift and northward tilt of the southernmost part of the peninsula. Radiocarbon dates on fossil shells from the three highest strandlines (Numa I-III terraces) yield an average uplift rate of 3.9 m/ka (Figure 6.25A). The lowest major strandline (Numa IV or Genroku terrace) at 5 m and a minor strandline at 1.0 m were formed by uplift during the great Kanto earthquakes of 1703 and 1923, respectively. The ages and vertical distances between successively higher pairs of dated strandlines yield an average uplift per event of about 6 m and an average recurrence interval of about 1.5 ka. More importantly, however, the uplift and age data appear to follow a FIGURE 6.24 Generalized tectonic map of Japan showing locations of emergent marine strandlines produced by historical coseismic uplift. O, Ogi Peninsula of Sado Island, 1802 earthquake (see Figure 6.13A for strandline profiles and Figures 6.15A and 6.24 for location); A, Awashima Island, 1964 Niigata earthquake; B, Boso Peninsula, 1703 and 1923 Kanto earthquakes (see Figure 6.25 for uplift history); M, Muroto Peninsula, 1707, 1855, and 1947 Nankaido earthquakes (see Figure 6.26A for uplift history; see Figure 6.16 for strandline and geodetic data); K, Kikai Jima, no historical coseismic strandlines, but see Figure 6.26B for Holocene uplift history.
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Active Tectonics: Studies in Geophysics FIGURE 6.25 Coseismic uplift, Boso Peninsula, Japan. A, Average uplift appears to be constant, and coseismic uplift events appear to follow a time-predictable pattern. B, Uplift pattern indicates that the next event (or sequence of events) will occur in A.D. 3000. Interearthquake subsidence (strain accumulation) from geodetic data is incorporated into prediction model. See Figure 6.24 for location and Figure 6.3A for profile of strandlines. Modified from Matsuda et al. (1978). time-predictable pattern. After being corrected for probable postearthquake subsidence (currently 1–2 mm/yr), the strandline data from the Boso Peninsula indicate that the next uplift event similar to the combined 1703 and 1923 events will occur in 0.8–1.3 ka (Figure 6.25B). Other workers (Nakata et al., 1980; and Shimazaki and Nakata, 1980; Yonekura and Shimazaki, 1980) also studied the emergent Holocene strandlines on the Boso Peninsula and reached similar conclusions. However, some of these workers (Yonekura and Shimazaki, 1980) suggested that the largest uplift events may have accompanied moderate-sized earthquakes on steep, secondary faults within the overthrust plate of the Sagami subduction zone and that the smallest uplift events may have accompanied larger earthquakes on the main but more gently dipping thrust fault. If this interpretation is correct, the 1703 Kanto earthquake may have been smaller than the 1923 event, even though the relative coseismic uplift would suggest otherwise. Interestingly, at Oiso, a densely populated suburb of Tokyo, about 70 km northwest of the Boso Peninsula, the sum of the uplifts that accompanied the 1703 and 1923 earthquakes is less than the uplift required to maintain the long-term uplift rate derived from older strandlines. This relationship suggests that the next uplift event in the Oiso area is overdue. Consequently, the Oiso area may currently be at greater seismic risk than the southern Boso Peninsula (Matsuda et al., 1978). At Murotsu harbor on the southern tip of the Muroto Peninsula on the island of Shikoku (Figure 6.24), the uplifts associated with the great Nankaido earthquakes of 1707 (1.8 m), 1855 (1.2 m), and 1947 (1.2 m) were determined by primitive but reasonably accurate harbor soundings and by a vertically displaced strandline defined by barnacles (Figures 6.16 and 6.26A) (Sawamura, 1953; Shimazaki and Nakata, 1980). These historical data yield a constant short-term uplift rate of 12.5 mm/yr, which is about six times higher than the FIGURE 6.26 Coseismic uplift on Muroto Peninsula and Kikai Jima, Japan. A, Historical coseismic strandlines at Murotsu Harbor. Uplift events appear to follow a time-predictable pattern and yield an average short-term uplift rate of 12.5 mm/yr, which is about six times higher than the long-term rate of 1.7–2.0 m/ka derived from 6-ka and 120-ka strandlines. If the short-term rate prevails, the next coseismic event should occur in the year 2040. If the long-term rate prevails, the next event (or sequence of closely spaced events) will not occur for another 3 ka. (See Figure 6.16 for more uplift data.) B, Radiocarbon-dated strandlines on Kikai Island, Japan, appear to follow a time-predictable pattern. If correct, a large coseismic uplift event should have occurred between 1400 and 1600. There is no geologic or historical evidence for such an event, so the area may be overdue for a large earthquake. Modified from Shimazaki and Nakata (1980). See Figure 6.24 for locations.
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Active Tectonics: Studies in Geophysics long-term rate of 1.7–2.0 m/ka derived from the 6-ka and 120-ka strandlines (Yoshikawa et al., 1964; Kanaya, 1978). If the exceptionally high uplift rate produced by historical coseismic events is maintained, the next large earthquake should occur in 2040 (Shimazaki and Nakata, 1980). However, if the long-term rate prevails, which seems more likely, the next event, or a series of closely spaced events, should not occur for another 3 ka. Tide-gauge and geodetic data from central Japan, which reveal the spatial and temporal patterns of interearthquake crustal deformation between the 1855 and 1947 events, indicate that because of the nonlinearity of strain buildup and the significant permanent deformation, simple recurrence calculations from strandline data may overestimate the true interval between major earthquakes by a factor of 2–3 in the Nankai area (Thatcher, 1984). On the other hand, the three closely spaced historical uplift events on the Muroto Peninsula and the two closely spaced events on the Boso Peninsula suggest that the large vertical spacing between emergent Holocene strandlines in these and possibly other areas may represent multiple, not single, earthquakes. If this interpretation is correct, earthquakes at a particular locality on a convergent plate boundary may occur in relatively tight clusters separated by considerable periods of time (1–3 ka), in which case the date of the next event would be difficult if not impossible to predict. On Kikai Island, the easternmost island in the Ryukyu archipelago of southern Japan, radiocarbon dates of fossil corals from four Holocene strandlines yield a uniform uplift rate of 1.8 m/ka over the past 6 ka and indicate that coseismic uplift followed a time-predictable pattern (Figures 6.24 and 6.26B) (Nakata et al., 1978, 1979; Shimazaki and Nakata, 1980). If the time-predictable model is correct, there should have been a large earthquake between A.D. 1400 and 1600. The lack of historical or geological evidence for an earthquake of this age suggests that Kikai Island, like the Oiso area near Tokyo, is overdue for a major earthquake (Shimazaki and Nakata, 1980). Two very different earthquake recurrence intervals are recorded by emergent Holocene strandlines on Sado and Awashima Islands in the intensely faulted area off the west coast of central Honshu, the largest Japanese island (Figure 6.24). Seven emergent Pleistocene and Holocene strandlines occur on the Ogi Peninsula, the southernmost tip of Sado Island (Figure 6.15A) (Ota et al., 1976). The lowest (2 m) strandline was formed by uplift during a major earthquake in 1802 (Figure 6.13A). This strandline and the next higher (4 m) 6-ka strandline are both tilted northward about 2.5×10−2, which indicates that the 1802 event was the only tilt event in the past 6 ka. The uniform difference of 2 m between these two strandlines is ascribed to a 2-m highstand of sea level at about 6 ka BP (Ota et al., 1976). A more likely explanation is that the area was uplifted uniformly after 6 ka BP but before the 1802 tilt event. In any case, dividing the coseismic tilt produced in 1802 into the long-term tilt rate derived from the 120-ka strandline yields a recurrence interval for seismic events of 5–9 ka, which is consistent with the >6-ka interval documented by the two lowest strandlines. At Awashima, which lies about 70 km northeast of Sado Island (Figure 6.24), tilted strandlines yield a much shorter recurrence interval (Nakamura et al., 1965). Awashima was uplifted a maximum of 1.8 m and tilted northward about 2.4×10−4 during the 1964 Niigata earthquake. The long-term tilt rate derived from the 82-ka strandline at 50–70 m above sea level is 4.3×10−4/ka. Dividing this tilt rate by the 1964 tilt yields an average recurrence interval of about 1.5 ka. The different earthquake recurrence intervals for the faults near the Ogi Peninsula and Awashima Island demonstrate that similar faults within the same tectonic province can have very different displacement histories. Alaska The narrow zone comprising the Aleutian archipelago and the south coast of mainland Alaska lies along the overthrust margin of the North American tectonic plate in the north Pacific Ocean. At numerous localities in this seismically active area, great interplate megathrust earthquakes are recorded by emergent Holocene and historical strandlines. On Middleton Island (Figure 6.18A) emergent Holocene strandlines record six coseismic uplift events over the past 4.5 ka (Figure 6.27A) (Plafker and Rubin, 1967, 1978). The lowest strandline at 3.5 m above sea level was produced by uplift associated with the great 1964 earthquake in southern Alaska. Radiocarbon dates on peat and wood from the wave-cut platforms of the five highest strandlines indicate that the time interval between uplift events increased gradually from about 0.5 ka to 1.4 ka and that the uplift rate decreased from 14 m/ka to 5.6 m/ka. If the island subsided slightly between major earthquakes, the uplift that accompanied each earthquake was greater than the vertical spacing between strandlines. It is not clear if the uplift events, which ranged from 7–12 m and averaged about 9 m, followed a displacement-predictable or a time-predictable pattern, but it is obvious that the 3.5-m uplift associated with the 1964 earthquake was not sufficiently large to maintain even the diminishing uplift rate (Figure 6.27A). In effect, at least half of the stress accumulated since the formation of the second lowest strandline at
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Active Tectonics: Studies in Geophysics FIGURE 6.27 Coseismic uplift, Gulf of Alaska. A, Radiocarbon-dated and historical coseismic strandlines on Middleton Island, Alaska. Average uplift rate is decreasing and interval between events is increasing while the size of each event has remained fairly constant. Average uplift is 0.95 m/ka. The 1964 coseismic uplift did not bring the island up to the long-term rate, which suggests that another event is due in the near future. Present sea level used as tectonic datum. Modified from Plafker and Rubin (1978). B, Radiocarbon-dated coseismic strandlines at Icy Cape (about 300 km east of Middleton Island on the Alaska mainland) yield a fairly constant uplift of about 1 m/ka. Uplift events appear to follow a time-predictable pattern. If this pattern continues, the next event is due in the near future. Modified from Plafker et al. (1981). See Figures 6.17C and 6.18A for 1964 coseismic strandline deformation. about 1.4 ka BP has yet to be released. Consequently, if the Holocene trend continues, coseismic uplift at least as large as the 1964 event should occur on Middleton Island sometime in the near future (Plafker and Rubin, 1978). The focal region of the 1964 Alaskan earthquake lies north and west of Middleton Island (Figure 6.18A), and a seismic gap lies to the east. Consequently, the next uplift event to affect the island may accompany a megathrust earthquake in the area of the seismic gap, possibly on the submarine Pamplona Fault (Plafker and Rubin, 1978). This interpretation is supported by data from Icy Cape, a broad sedimentary headland about 300 km east of Middleton Island on the Alaska mainland and within the area of the seismic gap. There, depositional strandlines dated at 4.9 ka, 2.4 ka, and 1.3 ka BP by radiocarbon techniques record at least three uplift events that followed a time-predictable pattern over the past 5 ka (Figure 6.27B) (Plafker et al., 1981). The lowest (1.3 ka) strandline may have been formed during the same seismic event that produced the second lowest strandline dated at 1.35 ka on Middleton Island (Figure 6.27A). If the 1.2–1.3-ka time span between strandlines at Icy Cape is the recurrence interval for the area of the present seismic gap, the next uplift event is overdue. Presumably, this event would make up some or all of the apparent deficiency in uplift on Middleton Island. An interesting aspect of the strandline sequence at Icy Cape is that the vertical distance between the 4.9-ka and 2.4-ka strandlines is 27 m (Figure 6.27B), which seems too large for a single uplift event. If uplift averaged 10–15 m per event, a 14-m uplift at about 3.8 ka BP may not have been recorded or its strandline may have been destroyed by subsequent wave erosion (Plafker et al., 1981). New Zealand The large southwest Pacific islands of New Zealand straddle the convergent boundary between the Australia-India tectonic plate to the west and the Pacific plate to the east. Northeast of the islands, the Australia-India plate is being thrust eastward over the Pacific plate, and southwest of the islands the relationship is reversed. The islands themselves are a manifestation of the uplift along the transition between the two opposing subduction zones on the same interplate boundary (Walcott, 1984). Within this complex structural area both interplate and intraplate earthquakes are common. At several localities, especially along the southeast coast of North Island, large earthquakes are recorded by emergent Holocene strandlines. At least five coseismic uplift events are recorded by six
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Active Tectonics: Studies in Geophysics bouldery beach ridges up to 27 m above sea level on the rapidly uplifting coastline at Turakirae Head, the southernmost tip of North Island (Figure 6.28) (Wellman, 1969; Stevens, 1973). The two lowest strandlines at 7 and 3 m above sea level, which were formed during historical earthquakes in 1460 and 1855, respectively, appear to record regional northwestward tilt across the entire southern tip of the island, whereas the four highest strandlines appear to record intense crustal warping between two local faults parallel to the offshore subduction zone (Wellman, 1969). However, all six strandlines may reflect regional tilt (Stevens, 1973). In either case, the five coseismic uplift events represented by these strandlines range from 2.5 to 9 m, and average 5.4 m (Figure 6.28). An assumption of constant uplift yields tentative ages for the four highest strandlines and also yields an average uplift rate of about 4 m/ka (see Figure 6.7) (Wellman, 1969). These graphically derived dates appear to indicate that uplift events follow a time-predictable pattern, which would indicate that the next event should occur in about 500 yr (Wellman, 1969). However, the pattern of uplift is inherent in the assumption of constant uplift—it is not independently demonstrated. Consequently, no firm estimate for the date of the next coseismic uplift event can be made. FIGURE 6.28 Coseismic Holocene strandlines at Turakirae Head, North Island, New Zealand. Highest (6 ka) strandline dated by graphical techniques (see Figures 6.5B and 6.7) assuming constant uplift rate (4 m/ka). Rate interpolated to estimate ages of lower strandlines. If uplift follows a time-predictable pattern, the next event should occur in 0.5 ka. It should be stressed, however, that these strandlines must be dated independently to demonstrate constant uplift and the time-predictable pattern. Modified from Wellman (1969). The often disparate and inconclusive estimates of past and future seismic events derived from sequences of emergent Holocene strandlines in the three areas discussed here (Japan, Alaska, and New Zealand) illustrate many of the difficulties encountered in interpreting even excellent historical and geological strandline data. However, in spite of these difficulties sequences of coseismic strandlines are among the most complete records of past earthquakes and in many coastal areas provide extremely valuable insights into the distribution of earthquakes in both space and time. FUTURE RESEARCH This brief review of coastal tectonics illustrates the breadth and scope of neotectonic deformation and seismic history that can be derived from the study of marine strandlines, which so conspicuously record the dynamic interaction between the fluctuating sea level and mobile tectonic plates along many of the world’s active coastlines. Without sea-level changes, discrete strandlines would not be produced and the long-term record of Pleistocene crustal movements in coastal areas would be extremely difficult if not impossible to extract from a relative sea-level record. Conversely, without vertical crustal movements, a detailed history of sea-level fluctuations would be virtually impossible to reconstruct. As our understanding of this complex interaction increases we discover new problems that require the reinterpretation of existing data and the acquisition of new, more precise information. It should be stressed, however, that even at our present level of understanding the uncertainties inherent in extracting past crustal movements and seismic histories from relative sea-level records are far outweighed by the wealth of useful information obtained. Following is a list of some of the most pressing problems and needs that must be addressed if we are to progress in the new and rapidly evolving field of coastal tectonics. Theoretical models relating sea-level changes and vertical crustal movements should be developed. General relationships are expressed by Eqs. (6.1) and (6.8) and Figures 6.5, 6.6, and 6.7. Theoretical models will provide a framework in which tectonic and sea-level data can be interpreted and evaluated. They will also provide graphical and mathematical means of correlating and dating strandlines. Models will also provide a coherent framework for compiling and comparing strandline data from different areas. The resolution of existing radiometric, chemical, and paleontologic dating techniques should be im-
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Active Tectonics: Studies in Geophysics proved and the ranges of these techniques extended. New dating techniques should be developed, especially for the period between 100–1000 ka BP. Both Pleistocene and Holocene (including historical) strandlines on many more coastlines throughout the world should be mapped and dated. Particular attention should be paid to precise ages and elevations. These data are needed to resolve potentially serious problems of geoidal distortion and regional hydro-isostatic crustal movements. If geoidal distortion is significant on a time scale of 10–100 ka, regional sea-level curves should be developed as tectonic datums. If geoidal distortion is relatively insignificant, its magnitude will help to establish uncertainty limits for rates of crustal deformation. In short, both Pleistocene and Holocene sea-level curves should be improved to provide more accurate estimates of tectonic deformation. 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Representative terms from entire chapter: