Bottom-Current Reworked Sands

Chapter 4

Bottom-Current Reworked Sands

4.1 INTRODUCTION The objective of this chapter is to discuss deep-water bottom currents and their deposits from oceanographic and sedimentological viewpoints. This is an expanded and updated version of earlier book chapters on this topic by Shanmugam (2006a, 2008c). The general term “bottom current” is used because it covers a variety of bottom currents of different origins, flow directions, and velocities (Shanmugam et al., 1993a, p. 1242). Southard and Stanley (1976) recognized five types of bottom currents at the shelf break based on their origin. These currents are generated by (1) thermohaline differences, (2) wind forces, (3) tidal forces, (4) internal waves, and (5) surface waves. In addition, tsunami-related traction currents have been speculated to occur in bathyal waters (Yamazaki et al., 1989), but the mechanics of such currents are not yet understood (Shanmugam, 2008b, 2011c). I have selected four types of deep-water bottom currents, namely (1) thermohaline-induced geostrophic bottom currents (i.e., contour currents), (2) wind-driven bottom currents, (3) deep-water tidal bottom currents, and (4) baroclinic currents (internal tides), for discussion.

4.2 SURFACE CURRENTS, DEEP-WATER MASSES, AND BOTTOM CURRENTS A sound knowledge of global ocean surface currents is critical for understanding ocean bottom currents. This is because surface currents and bottom currents are interrelated entities in the world’s oceans. For example, the three segments of the Southern Ocean, composed of (1) the upper surface currents, (2) the middle deep-water masses, and (3) the lower bottom currents, form a vertical continuum (Figure 4.1, see color plate). New Perspectives on Deep-water Sandstones. © 2012 Elsevier B.V. All rights reserved.

129

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Origin of deep-water masses

Vertical segments Polar front

Convergence

Ci

SC

Subantarctic surface water Subtropical surface water

rc cu um rre po nt lar

An

tar

cti

c

Antarctic surface water

ANT. intermediate water

DWM

Circumpolar deep water North Atlantic deep water

BC

Sinking of water mass

Antarctic bottom water

SC, Surface currents DWM, Deep-water masses BC, Bottom currents

FIGURE 4.1 A conceptual model of water mass bodies of the Southern Ocean showing three vertical segments, composed of the upper surface currents, the middle deep-water masses, and the lower bottom currents, forming a continuum (left). The origin of AABW by freezing of shelf waters (right). As a consequence, the increase in the density of cold saline (i.e., thermohaline) water triggers the sinking of the water mass down the continental slope and the spreading of the water masses to other parts of the ocean (see color plate). Modified after Hannes Grobe, April 7, 2000, http://en.wikipedia.org/wiki/File:Antarctic_bottom_water_hg.png. Accessed May 18, 2011.

FIGURE 4.2 Global circulation of ocean surface currents. From NOAA’s Ocean Motion (2011), http://oceanmotion.org/global-surface-currents.htm. Accessed March 20, 2011. Original image made by Michael Pidwirny.

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4.2.1 Surface Currents Basic information on ocean surface currents can be found in three websites maintained by (1) The Cooperative Institute for Marine and Atmospheric Studies (CIMAS, 2011), (2) the American Meteorological Society and National Oceanic and Atmospheric Administration (NOAA) (DataStreme Ocean, 2011), and (3) National Aeronautics and Space Administration (NASA) (Ocean Motion, 2011). Surface ocean currents are commonly winddriven entities that exhibit clockwise rotation in the Northern Hemisphere and counterclockwise rotation in the Southern Hemisphere (Figure 4.2, see color plate). Surface currents, which are mostly restricted to the upper 400 m of the ocean, make up nearly 10% of the water in the world’s oceans. In the Atlantic Ocean alone, CIMAS (2011) has identified 32 major ocean surface currents: 1. 2. 3. 4. 5. 6. 7. 8. 9. 10. 11. 12. 13. 14. 15. 16. 17. 18. 19. 20. 21. 22. 23. 24. 25. 26. 27. 28. 29. 30. 31. 32.

Agulhas (Gyory et al., 2004) Angola Antilles Azores Benguela Brazil Canary Caribbean System East Greenland East Iceland Florida (Section 7.4) Guiana Guinea Gulf Stream (Gyory et al., 2005) Irminger Labrador Loop Current Malvinas North Atlantic North Atlantic Drift North Brazil North Equatorial North Equatorial CC Norwegian/N Cape Portugal System Slope/Shelf edge Slope Jet South Atlantic South Equatorial System Spitsbergen Subtropical CC West Greenland

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4.2.2 Deep-water Masses The deep-water masses in the world’s oceans are caused by differences in temperature and salinity. When sea ice forms in the polar regions due to freezing of shelf waters, seawater experiences a concurrent increase in salinity due to salt rejection and a decrease in temperature. The increase in the density of cold saline (i.e., thermohaline) water directly beneath the ice triggers the sinking of the water mass down the continental slope and the spreading of the water masses to other parts of the ocean (Figure 4.1, see color plate). These are called thermohaline water masses. Stommel (1958) first developed the concept of the global circulation of thermohaline water masses and the vertical transformation of light surface waters into heavy deep-water masses in the oceans. Broecker (1991) presented a unifying concept of the global oceanic “conveyor belt” by linking the wind-driven surface circulation with the thermohaline-driven deep circulation regimes (Figure 4.3). The large-scale horizontal transport of water masses, which also sink and rise at select locations, is known as the “thermohaline circulation” or THC (Figure 4.3). The term THC, which refers to a driving mechanism by high-latitude cooling, is a physical concept and not a measurable quantity (Rahmstorf, 2006). A related concept is the “meridional

FIGURE 4.3 Map showing the global thermohaline circulation (THC) pattern of ocean currents. Red paths represent surface currents and blue paths represent deep-water currents. PSS, practical salinity scale. First published by Broecker (1991); later simplified by Rahmstorf (2002, 2006). See also Wunsch (2002) and Broecker (2006) for additional information. Map by Robert Simmon, NASA, http://earthobservatory.nasa.gov/Features/Paleoclimatology_Evidence/paleoclimatology_evidence_2.php. Accessed April 25, 2011.

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overturning circulation” (MOC), which refers to the north south flow as a function of latitude and depth. Although the terms THC and MOC are used synonymously (Zenk, 2008), there is no one-to-one relationship between the two (Rahmstorf, 2006). The global conveyor belt system in the North Atlantic originates near Greenland and Iceland where the sea-ice formation produces cold and salty North Atlantic Deep Water (NADW). The NADW sinks and flows southward along the continental slope of North and South America toward Antarctica where the water mass then flows eastward around the Antarctic continent. There, the NADW mixes with Antarctic waters (i.e., AABW (Antarctic Bottom Water) and AADW), forming the Antarctic Circumpolar water. These bottom waters gradually warm and mix with overlying waters as they flow northward. After rising to the surface in the Pacific, the surface waters flow through the many passages between the Indonesian islands into the Indian Ocean. Eventually they flow into the Agulhas Current (Figure 4.2, see color plate), which is the Indian Ocean boundary current that flows around southern Africa. Gordon (1985) measured the geostrophic volume transport of the Agulhas Current to be 67 Sv (1 Sv 5 1 3 106 m3 s 1). At the surface, the Agulhas Current can reach maximum speeds of 200 cm s 1 (Boebel et al., 1998). Schut et al. (2002) reported seismic evidence for bottom-current activity at the Agulhas Ridge. Examples of selected deep-water masses in various parts of the world’s oceans and their acronyms are: G G G G G G G G G G G G G G G G G G G G

AABW: Antarctic Bottom Water (Figure 4.1, see color plate) ABW: Arctic Bottom Water AAIW: Antarctic Intermediate Water (Brazilian margin) ACC: Antarctic Circumpolar Current (Antarctica) AW: Atlantic Water (Mediterranean Sea) BC: Brazil Current BICC: Brazil Intermediate Counter Current CDW or CPDW: Circumpolar Deep Water (Figure 4.1, see color plate) DGSRF: Deep Gulf Stream Return Flow DWBUC or DWBC: Deep Western Boundary Undercurrent LIW: Levantine Intermediate Water (Mediterranean Sea) MOW: Mediterranean Outflow Water MUC: Mediterranean Undercurrent NADW: North Atlantic Deep Water (Figure 4.1, see color plate) NAdDW: North Adriatic Dense Water NPDW: North Pacific Deep Water (Japan) NSDW: Norwegian Sea Deep Water SOW: Norwegian Sea Overflow Water SACW: South Atlantic Central Water (Brazilian margin) WBUC or WBU: Western Boundary Undercurrent

134 G G G

New Perspectives on Deep-water Sandstones

WDW: Warm Deep Water (Antarctica) WSBW: Weddell Sea Bottom Water (Antarctica) WSDW: Weddell Sea Deep Water (Antarctica).

4.2.3 Bottom Currents The deep-water component of these water masses that winnow, rework, and deposit sediment on the seafloor for a sustained period of time is called thermohaline-induced bottom currents (see Figure 4.1). In addition to thermohaline-induced bottom currents, there are three other major types, namely wind-driven, tide-driven, and internal tide-driven bottom currents.

4.3 BOTTOM CURRENTS VERSUS TURBIDITY CURRENTS In discussing deep-water bottom currents, it is imperative that we establish a clear distinction between bottom currents and turbidity currents. As defined in Chapter 1, turbidity current is a sediment flow with Newtonian rheology and turbulent state, in which sediment is supported by fluid turbulence and from which deposition occurs through suspension settling (Dott, 1963; Sanders, 1965; Middleton and Hampton, 1973; Shanmugam, 2000). Distinguishing deposits of deep-water bottom currents from those of turbidity currents has been a challenge (Bouma and Hollister, 1973; Stow, 1979; Mulder et al., 2009a). However, bottom currents and their deposits differ from turbidity currents and their deposits in the following respects: G

G

G

G

G

Bottom currents may occur on the shelf, slope, and in basinal environments, whereas turbidity currents are more common on the slope and basinal environments. Bottom currents are driven by thermohaline, wind, or tidal forces, whereas turbidity currents are always driven by sediment gravity. Bottom currents may flow parallel to the strike of the regional slope, may flow in circular motions (gyres) unrelated to the slope, or may flow up and down submarine canyons (tidal), whereas turbidity currents commonly flow downslope (Figure 4.4), though turbidity currents that flow parallel to the strike of the regional slope may occur due to local morphology, such as trench floors (Underwood and Bachman, 1982). Bottom currents have been documented by direct velocity measurements in the modern oceans, whereas no such direct velocity measurements exist for turbidity currents. Bottom currents persist for long periods of time and can develop equilibrium conditions, whereas turbidity currents are episodic (Kuenen and Migliorini, 1950) or surge-type events that fail to develop equilibrium conditions (Allen, 1973, 1985).

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FIGURE 4.4 Conceptual model showing the spatial relationship between downslope turbidity currents and along-slope bottom currents (i.e., contour currents). Wind-driven, tide-driven, and internal tide-driven bottom currents are not shown. After Shanmugam et al. (1993a), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

G

G

G

G

G

G

Bottom currents can exist without the presence of entrained sediment and, for this reason, they are termed “clear water currents” (Bouma and Hollister, 1973, p. 82), whereas turbidity currents, which are sedimentgravity flows, cannot exist without entrained sediment (Middleton and Hampton, 1973). Bottom currents show oscillating energy conditions, whereas turbidity currents exhibit waning energy conditions (Kuenen and Migliorini, 1950; Sanders, 1965). Bottom currents transport sand primarily by traction (i.e., bed-load movement by sliding, rolling, and saltation; Allen, 1984), whereas turbidity currents generally transport fine-grained sand and mud in suspension. Bottom-current sands are characterized by traction structures (e.g., parallel laminae, ripple laminae, and cross-beds) (Hsu, 1964; Shanmugam et al., 1993a; Mart
n-Chivelet et al., 2008), whereas normal grading is the norm in turbidites that are deposited by relatively catastrophic episodic events of waning energy (Kuenen and Migliorini, 1950). Bottom-current deposits generally exhibit sharp upper contacts (Hollister, 1967), whereas turbidites show gradational upper contacts. Bottom currents can result in well-sorted sand with good porosity and permeability because of reworking and winnowing away of mud (Shanmugam et al., 1993a), whereas turbidites are poorly sorted, commonly mud-rich deposits with low porosity and permeability (Pettijohn, 1957; Sanders and Friedman, 1997).

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4.4 GENETIC NOMENCLATURE The genetic term contourite was originally introduced for deposits of thermohaline-induced contour currents in the deep oceans (Hollister, 1967). Subsequently, the term has also been applied for deposits of wind-driven bottom currents in Lake Superior at water depths of 200 m (Johnson et al., 1980). These wind-driven bottom currents in Lake Superior did not originate as thermohaline-induced water masses in the polar regions. Therefore, these deposits in Lake Superior should not be classified as contourites. In this case, the original meaning of the term “contourite” has lost its oceanographic virtue. Other authors are also using the term “contourite” with a broader meaning (see review by Rebesco et al., 2008). This practice has its roots in turbidite research. For example, Mutti et al. (1999) expanded the meaning of the term “turbidite” to include “debrite”! In science, words should have clear and consistent meanings. In geology, however, this has not always been the case (Shanmugam, 2006b). The tradition of genetic nomenclature in sedimentary geology began with the introduction of the term “turbidite” for a deposit of a turbidity current in deepwater environments (Kuenen, 1957). Kuenen and Migliorini (1950, p. 99) and Kuenen (1967, p. 212) suggested that normal grading of a turbidite bed was a consequence of deposition from a single waning turbidity current. For a genetic term to succeed, (1) it must be based on sound fluid dynamic principles; (2) its usage must be accurate (relying on sedimentological description), precise (referring to a single process), and consistent (requiring a steady and a uniform application in time and space); and (3) it must imply a diagnostic flow behavior. Nonetheless, different authors have expanded the original meaning of the term “tubidite” to include deposits that are not turbidites. As a consequence, there is a plethora of turbidite nomenclature that include (1) atypical turbidites, (2) fluxoturbidites, (3) hemiturbidites, (4) high-concentration sandy turbidites, (5) megaturbidites, (6) problematica turbidites, (7) seismoturbidites, and (8) undaturbidites. And all these terms fail to reveal a clear flow behavior (Table 4.1). TABLE 4.1 Lexicon of Selected Genetic Terms Ending with “-ite” Genetic Terms

Comments (This Contribution)

Referencesa

Aeolianite

Implies the Aeolius (the god of the winds), not flow behavior

Sayles (1931); Bates and Jackson (1980)

Anastomosite

Implies river type, not flow behavior

Shanmugam (1984)

Atypical turbidite

Implies slumps, debris flows, and sand flows, not turbidity current

Stanley et al. (1978)

Braidite

Implies river type, not flow behavior

Shanmugam (1984) (Continued )

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Bottom-Current Reworked Sands

TABLE 4.1 (Continued) Genetic Terms

Comments (This Contribution)

Referencesa

Cascadite

Implies driving force (cascading), not depositional process

Gaudin et al. (2006)

Contourite

Implies current orientation, not flow behavior

Hollister (1967)

Debrite

Implies plastic debris flow

Pluenneke (1976)

Densite

Implies hybrid processes, not a single process

Gani (2004)

Diamictite

Implies pebbly mudstone, not flow (glacial) behavior

Flint et al. (1960)

Fluxoturbidite

A jargon with no discernible meaning (see Hsu¨, 1989)

Dzulynski et al. 1959)

Grainite

Implies grains, not flow behavior

Khvorova (1978)

Gravitite

Implies sediment gravity, not flow behavior

Natland (1967)

Gravite

Implies multiple processes, not a single process

Gani (2004)

Hemipelagite

Implies hemipelagic settling

Arrhenius (1963)

Hemiturbidite

Implies muddy turbidity current

Stow et al. (1990)

High-concentration sandy turbidite

Implies sandy debris flow, not turbidity current

Abreu et al. (2003)

Homogenite

Implies uniform grain size (ungraded mud), not flow behavior

Kastens and Cita (1981)

Hyperpycnite

Implies relative flow density, not flow behavior

Mulder et al. (2002)

Impactiteb

Implies impact by meteorite, not flow behavior

Sto¨ffler and Grieve (2003)

Injectiteb

Implies injection, not flow behavior

Vivas et al. (1988)

Interpretite

A spoof on genetic terms!

Davies (1997)

Meanderite

Implies river type, not flow behavior

Shanmugam (1984)

Megaturbidite

Implies debris flow, not turbidity current

Labaume et al. (1987)

Pelagite

Implies pelagic settling

Arrhenius (1963)

Seismiteb

Implies seismic shocks, not flow behavior

Seilacher (1984); Iqbaluddin (1978) (Continued )

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New Perspectives on Deep-water Sandstones

TABLE 4.1 (Continued) Genetic Terms

Comments (This Contribution)

Referencesa

Seismoturbidite

Implies mass flow (debris flow), not turbidity current

Mutti et al. (1984a)

Suspensite

Implies suspension settling

Lisitsyn (1986)

Tectonite

Implies tectonic deformation, not flow behavior

Turner and Weiss (1963)

Tempestite

Implies multiple processes, not a single process

Ager (1974)

Tidalite

Implies deposition from tidal currents

Klein (1971, 1998)

Tillite

Implies pebbly mudstone, not flow (glacial) behavior

Harland et al. (1966)

Tractionite

Implies traction deposition by bottom current

Natland (1967)

Tsunamite

Implies multiple processes, not a single process

Gong-Yiming (1988)

Turbidite

Implies turbulent turbidity current

Kuenen (1957)

Undaturbidite

A jargon with no discernible meaning

Rizzini and Passega (1964)

Unifite

Implies grain size (ungraded mud), not flow behavior

Feldhausen et al. (1981); Stanley (1981)

Winnowite

Implies winnowing action of bottom current

Shanmugam and Moiola (1982)

Source: Modified after Shanmugam (2006b). a References include those that introduced the term, used the term early, or considered appropriate. b Unrelated to depositional processes. Notes: Items in bold are related to bottom-current reworking. Most genetic terms are obsolete in depositional process sedimentology.

Like turbidites, genetic terms of bottom-current reworked sands (BCRS) also fail to divulge diagnostic flow behavior. Selected examples are as follows: 1. The term contourite emphasizes current orientation with respect to bathymetric contours (Hollister, 1967), not the flow behavior. 2. The term laminite represents sedimentary structure (i.e., lamina) (Lombard, 1963), not the flow behavior. 3. The term tractionite implies traction deposition from bottom currents (Natland, 1967), but traction deposition has also been attributed to turbidity currents (Middleton and Hampton, 1973).

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4. The term tidalite implies alternating units of traction and suspension deposition from shallow-water tidal currents (Klein, 1971). Deepwater tidal bottom currents can also develop alternating units of traction and suspension deposition (Klein, 1975; Shanmugam, 2003). Therefore, this term can be used to convey a specific flow behavior. 5. The term winnowite implies winnowing action of bottom currents (Shanmugam and Moiola, 1982), but this process of reworking takes place after the primary deposition. 6. The term tsunamite represents tsunami-induced tractive bottom-current reworked sediment (Yamazaki et al., 1989), but the term tsunami is not a self-defining expression of a single depositional process (Shanmugam, 2006b). In summary, the genetic terms contourite and tidalite can be used to convey a specific bottom-current process, but most other genetic terms are ineffective.

4.5 THERMOHALINE-INDUCED GEOSTROPHIC BOTTOM CURRENTS 4.5.1 Antarctic and Arctic Bottom Currents I have already introduced this topic in Section 4.2. Concepts of deep-marine bottom currents in modern oceans became popular when Heezen et al. (1966) reported deep-water masses that flow along the ocean floor. An example of such bottom currents is the AABW (Figure 4.1, see color plate). AABW was first identified by Brennecke (1921) in the northwest corner of the Weddell Sea in the Antarctic region. The origin of the AABW was attributed to the formation of ice from surface freezing over the Antarctic continental shelves (Figure 4.1, see color plate). The Western Boundary Undercurrent (WBUC or WBU), the Arctic counterpart to AABW, originates as a cold dense-water mass from the Norwegian Sea off Greenland (Worthington and Volkman, 1965). It flows southward along the western margin of the North Atlantic. These thermohaline currents tend to flow parallel to the slope, that is, along the slope at right angle to downslope flowing gravity-induced turbidity currents (Figure 4.4). The WBUC is deflected in the Northern Hemisphere to the west as a result of the Coriolis force. Because of its tendency to flow parallel to bathymetric contours, the WBUC is known as a contour current (Heezen et al., 1966). These currents are also called geostrophic contour currents, because they strike a balance between the Coriolis and the gravity forces. Although contour currents and their deposits (contourites) have received a skewed emphasis in the published literature, there are no quantitative data to prove that contour currents and their deposits are geologically more important than the other three types (wind, tide, and internal tide). As noted earlier, the THC is not a measurable quantity (Rahmstorf, 2006).

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4.5.2 Current Velocity Maximum current velocities of bottom currents in different parts of the world’s oceans are summarized in Table 4.2. Measured current velocities usually range from 1 to 20 cm s 1 (Hollister and Heezen, 1972); however, exceptionally strong, near-bottom currents with maximum velocities of up to 300 cm s21 were recorded in the Strait of Gibraltar (Gonthier et al., 1984) (Table 4.2). Bottom-current velocities of 73 cm s21 were measured at a water

TABLE 4.2 Maximum Current Velocities of Bottom Currents in Different Parts of the World’s Oceans Study Area

Depth (m)

Maximum Current Velocity (cm s21)

Straits of Gibraltar (Gonthier et al., 1984)

400 1,400

300

Upper slope. Offshore Brazil, Equatorial Atlantic (Viana et al., 1998)

200

300

Gulf of Mexico, Loop Current (Cooper et al., 1990)

100

204

Green Canyon 166 area, Gulf of Mexico. Drilling operations were temporarily suspended in August of 1989 because of high-current velocities that reached 153 cm s21 (Koch et al., 1991).

45

153

Faeroe Bank Channel, North Atlantic (Crease, 1965)

760

109

Rise, Off Nova Scotia, North Atlantic (Richardson et al., 1981)

5,000

73

Base of North American Continental Rise (Bulfinch and Ledbetter, 1983/84)

5,022

73

Trench, Ryukyu Trench, Japan (Tsuji, 1993)

340

51

Samoan Passage, Western South Pacific (Hollister et al., 1974)

50

Hebrides Slope, North Atlantic (Howe and Humphrey, 1995)

403 468

48

Faeroe-Shetland Channel, North Atlantic (Akhurst, 1991)

900

33

Rise, near Hatteras Canyon, North Atlantic (Rowe, 1971) Carnegie Ridge, Eastern Equatorial Pacific (Lonsdale and Malfait, 1974)

33 1,000 2,000

.30

(Continued )

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TABLE 4.2 (Continued) Study Area SE of Iceland, North Atlantic (Steele et al., 1962)

Depth (m) 2,100 slope

Argentine Basin, Western South Atlantic (Ewing et al., 1971)

Maximum Current Velocity (cm s21) 30 30

Amirante Passage, Western Indian Ocean (Johnson and Damuth, 1979)

4,000 4,600

30

Rise, Off New England, North Atlantic (Zimmerman, 1971)

3,000 5,000

26.5

Blake-Bahama Outer Ridge, North Atlantic (Amos et al., 1971)

4,300 5,200

26

Off North Carolina, North Atlantic (Rowe and Menzies, 1968)

1,500 4,000

25

Off Cape Cod, North Atlantic (Volkman, 1962)

10 3,200

21.5

Off Cape Hatteras, North Atlantic (Barrett, 1965)

21

Greater Antilles Outer Ridge, North Atlantic (Tucholke et al., 1973)

5,300 5,800

20

Off Blake Plateau, North Atlantic (Swallow and Worthington, 1961)

3,300 3,500

20

Tonga Trench and vicinity, Western South Pacific (Reid, 1969)

.4,800

19

Western North Atlantic (Wust, 1950)

2,000 3,000

17

West Bermuda Rise, North Atlantic (Knauss, 1965)

5,200

17

Scotia Ridge, Antarctic Circumpolar Current, Antarctica (Zenk, 1981)

3,008

17a

Greenland-Iceland-Faeroes Ridge, North Atlantic (Worthington and Volkman, 1965)

2,000 3,000

12

Antillean-Caribbean Basin (outer), North Atlantic (Wust, 1963)

4,000 8,000

10

a

1-year vector averaged speed.

depth of 5 km on the lower continental rise off Nova Scotia (Richardson et al., 1981). Heezen and Hollister (1971) suggested that at extremely high bottom velocities of over 100 cm s21, relict pockets of sand and gravel may

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occur on the barren seafloor. According to Bulfinch and Ledbetter (1983/ 1984), a Deep Western Boundary Undercurrent (DWBUC) flows south along the North American continental slope and rise between 1,000 m and 5,000 m. The DWBUC has a 300-km wide high-velocity zone, with a maximum measured velocity of 73 cm s21, which winnows both fine and very fine silt, and results in deposition of medium and coarse silt.

4.5.3 Sedimentological Criteria Stow and Lovell (1979) have recognized two types of contourites, namely muddy contourites and sandy contourites. In this chapter, I focus on sandy contourites because of their relevance in petroleum reservoirs. The following general features of thermohaline-induced current deposits (i.e., sandy contourites) have been discussed by Hubert (1964), Hollister (1967), Hollister and Heezen (1972), Bouma and Hollister (1973), Unrug (1977), Stow and Lovell (1979), Lovell and Stow (1981), Shanmugam (2000), and Ito (2002): G G

G

Fine-grained sand and silt. Thin-bedded to laminated sand (usually less than 5 cm) associated with deep-marine mud (core and outcrop) (Figure 4.5). Rhythmic occurrence of sand and mud layers (core and outcrop) (Figure 4.6).

Lenticular layers

Ripple cross-laminae

Sharp upper contact

2 cm

FIGURE 4.5 Core photograph showing well-sorted fine-grained sand and silt layers (light gray) with interbedded mud layers (dark gray). Note sand layers with sharp upper contacts, internal ripple cross-laminae, and mud-offshoots. Also note lenticular nature of some sand layers. Pleistocene, continental rise off Georges Bank, Vema 18 374, 710 cm, water depth 4,756 m. After Hollister (1967, his Figure VI-1, p. 208) and Bouma and Hollister (1973), reproduced with permission from SEPM.

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G G G G

143

Bottom-Current Reworked Sands

Inverse size grading (core and outcrop) (Figure 4.6). Sharp to gradational bottom contacts (core and outcrop). Sharp, nonerosional, upper contacts (core and outcrop) (Figure 4.5). Horizontal laminae (core and outcrop) (Figures 4.6 4.8).

Inverse grading

FIGURE 4.6 Core photograph showing rhythmic layers of sand and mud, inverse grading, and sharp upper contacts of sand layers (arrow), interpreted as bottom-current reworked sands. Paleocene, North Sea. After Shanmugam (2008c), with permission from Elsevier.

Horizontal laminae

FIGURE 4.7 Core photograph showing horizontal laminae. Eocene, North Sea.

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Horizontal laminae

FIGURE 4.8 Outcrop photograph showing horizontal lamination. Calcareous Contourites, Maastrichtian, Caravaca, Spain. After Martin-Chivelet et al. (2008), with permission from Elsevier.

Cross laminae

5 cm

FIGURE 4.9 Outcrop photograph showing low-angle cross-lamination. Calcareous Contourites, Maastrichtian, Caravaca, Spain. After Martin-Chivelet et al. (2008), with permission from Elsevier. G G G G

Low-angle cross-laminae (core and outcrop) (Figure 4.9). Cross-bedding (core and outcrop) (Figure 4.10). Ripple cross-laminae (core and outcrop) (Figure 4.11). Cross-laminae accentuated by heavy mineral placers (Bouma and Hollister, 1973).

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Bottom-Current Reworked Sands

Cross-bedding

FIGURE 4.10 Outcrop photograph showing cross-bedding. Sandy contourites, Brushy Canyon Member, Permian, Delaware Basin, West Texas. Photo courtesy of E. Mutti. See Mutti (1992) for more details.

Ripple cross-laminae

FIGURE 4.11 Core photograph showing current ripples in fine-grained sand. Paleocene, North Sea.

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New Perspectives on Deep-water Sandstones

Rippled sand (Bouma Tc?) Erosional contact

Massive sand (Bouma Ta?)

Sand unit with ripple lamination near top

FIGURE 4.12 Core photograph showing basal massive sand division and upper rippled sand division, mimicking Bouma Ta and Tc divisions, respectively. Note internal erosional contact. The ideal “Bouma Sequence” (Appendix A) does not contain internal erosional contacts. Paleocene, North Sea. G

G G

Massive sand division overlain by rippled sand division, mimicking Bouma Ta and Tc divisions but with an internal erosional contact (core and outcrop) (Figure 4.12). The ideal “Bouma Sequence” does not contain internal erosional contacts. Also, the problem remains how we can explain deep-water units that show a partial Bouma-like units composed of a basal massive division and an upper parallel or ripple laminated division (see “The Bouma Sequence” in Appendix A for details). In areas where both downslope sandy debris flows and along-slope bottom currents operate concurrently (Figure 4.13), the reworking of the tops of sandy debrites by bottom currents may be expected. Such a scenario, common on the U.S. Atlantic margin, could generate a basal massive sand division and an upper reworked division, mimicking a partial Bouma Sequence. The reworking of deep-water sands by bottom currents has been suggested by other researchers as well (e.g., Stanley, 1993; Ito, 2002). Lenticular bedding or starved ripples (core and outcrop) (Figure 4.14). Well-sorted sand; little depositional mud matrix (clean sand) (core and outcrop).

No single criterion by itself is unique for distinguishing deposits of thermohaline-induced bottom currents. Although many of the criteria listed

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Bottom currents Muddy turbidity current Sandy debris flow Hydroplaning Muddy debris flow Turbidity current Hemipelagic setting

FIGURE 4.13 Conceptual model showing reworking the tops of downslope sandy debris flows by along-slope bottom currents. Such complex deposits would generate a sandy unit with a basal massive division and upper reworked divisions with traction structures (ripple laminae), mimicking the “Bouma Sequence” Ta and Tc divisions. After Shanmugam (2006a), with permission from Elsevier.

FIGURE 4.14 Core photograph showing lenticular laminae. Middle Oligocene, DSDP Leg 28, Site 268. Antarctica. From Piper and Brisco (1975).

above can be attributed to processes other than bottom-current reworking, the association of several of the above criteria in a given deep water example, along with the knowledge of the regional depositional setting, greatly enhances the chance of recognizing bottom-current reworked facies. But still it is impossible to establish that a given sedimentary structure in the rock record was originated by contour-following thermohaline currents, without establishing the paleo-water circulation pattern independently. Therefore, the general term “bottom-current reworked sands” is appropriate in many cases.

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7363 ft

Concave-up laminae (Wave ripple)

Convex-up laminae (HCS?)

FIGURE 4.15 Core photograph showing core-scale convex-up laminae (HCS, hummocky cross-stratification?) and concave-up laminae (wave ripples) in fine-grained sand. Eocene, North Sea.

During the course of my investigation of cores from the North Sea and Norwegian Sea, I have also observed the following features in deep-water strata: G

G G G

Convex-up (HCS, hummocky cross-stratification?) and concave-up forms (wave ripples) (core and outcrop) (Figure 4.15). Aggradational (wave) ripples (core and outcrop) (Figure 4.16). Diffused double mud layers (core and outcrop) (Figure 4.17). Flaser bedding (core and outcrop) (Figure 4.18).

Although the paleogeographic location of these samples indicates the possible influence of paleo-contour currents in the region, these features also suggest the influence of wave- and tide-induced bottom currents. The problem here is that the origin of HCS is controversial: a. HCS has traditionally been considered as evidence for storm-wave deposition by oscillatory flows in shallow-marine (,200 m water depth) environments (Harms et al., 1975). b. Duke et al. (1991) proposed the origin of HCS by combined flows (i.e., combination of unidirectional flows and oscillatory flows).

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Aggradational (wave) ripple

FIGURE 4.16 Core photograph showing aggradational ripples (HCS, hummocky crossstratification?) in fine-grained sand. Eocene, North Sea.

Double mud layers

FIGURE 4.17 Core photograph showing double mud layers (arrows) in fine-grained sand. Lysing Formation, Cretaceous, offshore Norway.

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Mud-draped ripples

FIGURE 4.18 Core photograph showing mud-draped ripples. Agat Formation, Cretaceous, offshore Norway.

c. Mutti (2009) suggested the origin of HCS by hyperpycnal flows in deltafront deposits. d. Mulder et al. (2009b) explained the origin of HCS-like structures as deep-sea antidune stratification formed by turbidity-current standing waves, relating to Kelvin Helmholtz instability occurring between highdensity basal layer and low-density upper layer in stratified flows (Prave and Duke, 1990). Higgs (2010) disputes this deep-sea origin of HCS. Furthermore, the concept of stratified layers in high-density turbidity currents is itself in dispute (Shanmugam, 1996a). These examples reveal the practical challenges that exist in classifying bottom-current deposits into a particular origin (i.e., thermohaline, wind, and tide). Chances are that more than one type of bottom current may operate at a given time in a given place. Therefore, a complex mix of sedimentary features is expected from deep-water deposition. For this reason, genetic facies models for bottom-current deposits are inappropriate.

4.5.4 Problematic Contourite Facies Model A contourite facies model, first proposed by Fauge`res et al. (1984), consists of a basal inversely graded unit overlain by a normally graded unit (Gonthier et al., 1984; Stow et al., 1998). Stow and Fauge`res (2008, their Figure 13.9)

Lithology and structure 4

Mean grain size 8 16 32 64 µm

C4

Mottled silt and mud

C3

Sandy silt

C2

Mottled silt and mud

C1

Mud

Laminated (diffuse) ± Cross-laminated silt Bioturbated Silt mottles and lenses irregular Horizontal alignment Bioturbated Massive irregular sandy pockets Bioturbated Contacts sharp to gradational ± hiatuses Silt mottles, lenses and irregular layers Bioturbated Contacts variable Indistinct lamination Bioturbated

Bioturbation

Discontinuous silt lenses

Gradational contact

Sharp (irregular) contact

Maximum velocity Negatively graded

Mud

Positively graded

Bioturbated

C5

ase ecre city d lo e V

Proposed divisions

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Botto velocit m current y incr ease

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10 (cm) 0

FIGURE 4.19 Contourite facies model with five divisions. After Stow and Fauge`res (2008), with permission from Elsevier.

recently revised the model and proposed a predictable vertical sequence with five bioturbated divisions (Figure 4.19). Major problems with this model are as follows. 1. Analogous to the five divisions (Ta, Tb, Tc, Td, and Te) of the turbidite facies model (i.e., the “Bouma Sequence,” see Appendix A), Stow and Fauge`res (2008) have introduced five divisions (C1, C2, C3, C4, and C5) for their contourite facies model. Stow and Fauge`res do not explain the fluid mechanics behind the origin of each of the five divisions. In defense of their facies model, Stow and Fauge`res (2008, p. 240) have argued that process models derived from ancient strata are less reliable in comparison to their contourite model derived from modern sediments. Their argument is specious because the “Bouma Sequence,” which they have used as the analogy for their model, was derived strictly from a study of ancient strata (Bouma, 1962). 2. No one has reproduced the complete “Bouma Sequence” in laboratory flume experiments. Nor has anyone replicated the complete contourite sequence in experiments. 3. The complete Bouma Sequence with all five divisions has never been documented in sediments on the modern seafloor.

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4. There is no explanation as to how one could infer the current orientation (i.e., contour-following currents) from vertical changes in grain size (i.e., inverse-to-normal grading) advocated in the model (Figure 4.19). Without the critical information on contour-following currents, the term “contourite” is meaningless for the model. 5. Inverse-to-normal grading, common in hyperpycnites (Mulder et al., 2003) (see also Fauge`res and Mulder, 2011, their Figure 2.12C), is not unique to contourites. 6. Stow and Fauge`res (2008, p. 235) state that “. . .well-defined sedimentary structures are generally absent, in part because they have been thoroughly destroyed by bioturbation.” This logic goes against the founding principle of process sedimentology in which primary sedimentary structures are the sole foundation for interpreting depositional mechanics (Sanders, 1963). The danger here is that one could also use the logic of Stow and Fauge`res (2008) for interpreting any bioturbated deep-water unit with inverse-to-normal grading as “turbidite” as well. 7. Bioturbation is a diagnostic criterion of the contourite facies model. This is based on the belief that active contour currents would increase the oxygen concentration of the water mass, and thereby would increase the activity by benthic organisms. Tucholke et al. (1985), on the other hand, suggested that the degree of preservation of bioturbation is a function of bottom-current intensity; strong bottom currents do not favor preservation of biogenic structures. Bioturbated mud in the deep sea is equally abundant in areas that are unaffected by contour currents. Even if bioturbation were prevalent in areas of contour currents, it would not directly reveal anything unique about current orientation (i.e., contour-following currents). The bioturbation criterion for “contourites” is seriously flawed because ancient deep-water turbidites (e.g., in the Late Cretaceous Point Loma Formation near San Diego, California) are extensively bioturbated and contain the trace fossil Ophiomorpha (Nilsen and Abbott, 1979). In contrast, convincing cases of “contourites” without bioturbation have been documented in the rock record (Dalrymple and Narbonne, 1996). 8. Recently, Fauge`res and Mulder (2011) have published their most up-todate views on contour currents and contourites. In this review, they have advocated the contourite facies model originally published by Fauge`res et al. (1984). One wonders as to why Fauge`res and Mulder (2011) have reverted back to the original model published in 1984 without the five internal divisions, ignoring the most recent version by Stow and Fauge`res (2008) with five internal divisions. It is puzzling because Fauge`res is the coauthor of both publications. 9. Both the original and the revised contourite facies models are fundamentally flawed. This is because none of the three criteria, namely (a) lack of primary structures, (b) inverse-to normal-grading, and (c) bioturbation, reveals anything unique about contour-following bottom currents in order

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to qualify the deposit as contourites, which is the basic tenet of the “contourite” model. In my view, genetic turbidite and contourite facies models are nothing more than a “groupthink” and a step backward in the pragmatic science of process sedimentology.

4.6 WIND-DRIVEN BOTTOM CURRENTS 4.6.1 The Gulf Stream The Gulf Stream is a powerful, warm, and swift Atlantic Ocean current that originates at the tip of Florida, and follows the eastern coastlines of the United States and Newfoundland before crossing the Atlantic Ocean (Figure 4.20). The Gulf Stream proper is a western-intensified current, largely driven by wind stress (Wunsch, 2002). The Gulf Stream is known to break off, forming separate eddies (Figure 4.20). The Gulf Stream is

Atlantic Ocean

United States

Eddy Gulf Stream

FIGURE 4.20 Moderate Resolution Imaging Spectroradiometer (MODIS) image showing the Atlantic Ocean’s Gulf Stream on May 2, 2001. The false colors in the image represent “brightness temperature” observed at the top of the atmosphere. The brightness temperature values represent heat radiation from a combination of the sea surface and overlying moist atmosphere. The red pixels in this image show the warmer areas (approaching 250 C), green pixels are intermediate values (12 130 C), and blue pixels are relatively low values (less than 100 C). Note the eddy of the Gulf Stream. Credit, NASA and Liam Gumley, http://eoimages.gsfc.nasa.gov/ve/ 1722/modis_brighttemp_glfstr_lrg.jpg. Accessed March 20, 2011.

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100 200 km wide (CIMAS, 2011) and extends to depths of 1,200 m (3,900 ft). The current velocity is fastest near the surface, with a maximum velocity of about 250 cm s21 (9 km hr21), confined to the upper 200 m of the water column. Hollister (1967) suggested that the currents associated with the Gulf Stream may reach the deep-sea floor in the Atlantic. Such bottom currents have been reported beneath the Gulf Stream Gyre at a depth of nearly 4 km in the northern Bermuda Rise (Laine, 1978). Laine and Hollister (1981) suggest that the Deep Gulf Stream Return Flow (DGSRF) entrains suspended sediment in a deep gyre and may be responsible for the deposition at the base of the continental rise.

4.6.2 The Loop Current The Loop Current in the eastern Gulf of Mexico is a wind-driven surface current (Figure 4.21, see color plate), and it is genetically linked to the Gulf

United States

Gulf Stream

Loop Current

Florida Current

FIGURE 4.21 Sea surface temperature (SST) image showing the Loop Current in the Gulf of Mexico and the axis of the Gulf Stream in the Atlantic Ocean along the U.S. Continental margin on March 12, 2011. Darker orange to red color enhancement represents temperatures in the upper 70
s F. Image credit: NOAA’s Cooperative Institute for Meteorological Satellite Studies (CIMSS), University of Wisconsin  Madison, USA, http://cimss.ssec.wisc.edu/goes/blog/wpcontent/uploads/2011/03/MODIS_SST_20110312_1615_largescale.png. Accessed March 29, 2011.

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155

FIGURE 4.22 Present circulation pattern of the Loop Current (After Neumann and Pierson, 1966; Nowlin, 1972; and Mullins et al., 1987). This wind-driven surface current is considered to be affecting the sea bottom (Pequegnat, 1972). Note the eddies detached from the Loop Current in the Ewing Bank area. After Shanmugam et al. (1993a), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

Stream in the Atlantic Ocean. The Loop Current enters the Gulf of Mexico through the Yucatan Strait as the Yucatan Current; it then flows in a clockwise loop in the eastern Gulf as the Loop Current, and exits via the Florida Strait as the Florida Current (Neumann and Pierson, 1966; Nowlin, 1972; Mullins et al., 1987). Finally, this current merges with the Antilles Current to form the Gulf Stream (Figure 4.22). The Loop Current also propagates eddies into the north-central Gulf of Mexico, where the Ewing Bank area is located, a case study used in this book (Figure 4.23).

4.6.3 Current Velocity Velocities in eddies that have detached from the Loop Current have been recorded as high as 200 cm s21 at a depth of 100 m (Cooper et al., 1990). The Loop Current and related eddies pose significant problems for deepwater drilling (Koch et al., 1991). For example, drilling operations in the

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FIGURE 4.23 Location map of the Ewing Bank and adjacent areas in the Northern Gulf of Mexico. Solid dots show locations of cores. After Shanmugam et al. (1993a), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

Green Canyon 166 area were temporarily suspended in August of 1989 because of high-current velocities that reached nearly 150 cm s21 at a depth of 45 m, and 50 cm s21 at a depth of 250 m. These intense bottom currents affect the ability of a drilling rig to hold station over a wellhead (Koch et al., 1991). Current-velocity measurements, bottom photographs, high-resolution seismic records, and GLORIA side-scan sonar records indicate that the Loop Current influences the seafloor at least periodically in the Gulf of Mexico (Pequegnat, 1972). Computed flow velocities of the Loop Current vary from nearly 100 cm s21 at the sea surface to more than 25 cm s21 at 500 m water depth (Nowlin and Hubert, 1972). This high surface velocity suggests a wind-driven origin for these currents. Flow velocities measured using a current meter reach up to 19 cm s21 at a depth of 3,286 m (Pequegnat, 1972). Kenyon et al. (2002b) reported 25 cm s21 current velocity measured 25 m above the seafloor. Such currents are capable of reworking fine-grained sand on the seafloor. Current ripples, composed of sand at a depth of 3,091 m on the seafloor (Figure 4.24), are the clear evidence of deep bottom-current activity in the Gulf of Mexico today (Pequegnat, 1972). These current ripples are draped by thin layers of mud. If these mud drapes on sand ripples were preserved in the rock record, they would be termed “mud-offshoots.”

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NE-16 3091 m

50 cm

FIGURE 4.24 Undersea photograph showing possible mud-draped (arrow) current ripples at 3,091 m water depth in the Gulf of Mexico. Similar mud drapes may explain the origin of mudoffshoots observed in the core (see Figure 4.36). A current measuring nearly 18 cm s21 was recorded on the day this photograph was taken. Current flow was from upper left to lower right. Alaminos Cruise 69-A-13, St.35. From Pequegnat (1972), reproduced with permission from Gulf Publishing Company.

4.6.4 The Loop Current Tropical Cyclone Connection The Loop Current is influential in causing sediment failures in the northern Gulf of Mexico through enhancing wind velocity of tropical cyclones. For example, there are empirical data to show that the wind speeds of Hurricane Katrina increased dramatically as it passed through the warm waters of the Loop Current toward the Gulf Coast in late August in 2005 (Figure 4.25). This increased velocity of hurricanes has implications for triggering sediment failures (Chapter 5). On the outer continental shelf (OCS) of the Gulf of Mexico, for example, the 2005 category 5 Hurricane Katrina destroyed 46 petroleum platforms and damaged 20 others (MMS, 2006). Katrina-induced mudflows damaged at least six pipelines (Alvarado, 2006).

4.6.5 Sedimentological Criteria Deposits of the Loop Current have been interpreted in the cores from the Ewing Bank 826 Field, Plio-Pleistocene, Gulf of Mexico. The Ewing Bank Block 826 Field is located nearly 100 km off the Louisiana coast in the

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United States

Cat. 4

Track of Hurricane Katrina

Cat. 5 Cat. 4 Cat. 3

Gulf of Mexico

Cat. 2

Loop Current

FIGURE 4.25 Image of sea surface height (cm) showing that the warm waters of the Loop Current (red) were 50 to 75 cm higher than the surrounding water when Hurricane Katrina passed through the Loop Current during August 26, 27, 28, and 29, 2005. Note that the wind speeds (mi hr21) of Hurricane Katrina increased dramatically as it passed over the warm waters of the Loop Current toward the Gulf Coast. Hurricane Katrina’s wind speed is highlighted by Saffir Simpson Scale categories 2 5 (see “Tropical Cyclone” in Appendix A). The image was produced by a University of Colorado at Boulder team, and processed at CU-Boulder’s Colorado Center for Astrodynamics Research (CCAR). Image credit, http://www.colorado.edu/ news/releases/2005/358.html. Accessed March 31, 2011.

northern Gulf of Mexico (Figure 4.23). It contains hydrocarbon-producing clastic reservoir sands that have been interpreted as BCRS (Shanmugam et al., 1993a, 1993b). Cores from the Ewing Bank and adjacent areas exhibit the following features: G G

G

G

Predominantly fine-grained sand and silt (core and outcrop). Thin-bedded to laminated sand (usually less than 5 cm) intercalated with deep-water mud (core and outcrop) (Figure 4.26). Sharp (nonerosional) upper contacts of sand layers (core and outcrop) (Figure 4.26). Sharp to gradational bottom contacts (core and outcrop).

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159

FIGURE 4.26 Core photograph showing discrete thin sand layers with sharp upper contacts (top arrow). Traction structures include horizontal laminae, low-angle cross-laminae (middle arrow), and starved ripples (bottom arrow). Dip of cross-laminae to the right suggests current from left to right. Note rhythmic occurrence of sand and mud layers. Middle Pleistocene, Gulf of Mexico. After Shanmugam et al. (1993a), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use. G

G

G G

G G G

G G G

G G

Horizontal lamination and low-angle cross-lamination (core and outcrop) (Figure 4.26). Lenticular bedding or starved ripples at core scale (core and outcrop) (Figure 4.26). Rhythmic sand and mud layers (core and outcrop) (Figure 4.27). Numerous layers (50 or more per 1 m of core) (core and outcrop) (Figure 4.27). Internal erosional surfaces (core and outcrop) (Figure 4.28). Erosional truncation surfaces (core and outcrop) (Figure 4.29). Bidirectional cross-lamination (core and outcrop) (Figure 4.30). This feature is typical of tidal currents. Megascopic inverse size grading (core and outcrop) (Figure 4.31). Microscopic inverse size grading (core and outcrop) (Figure 4.32). Current ripples with preserved crest or with eroded crest (core and outcrop) (Figure 4.33). Ripple forms with curved bases (core and outcrop) (Figure 4.33). Ripple cross-lamination in discrete sand units (core and outcrop) (Figure 4.34). This kind of discrete sand units composed solely of traction structures can be better explained by bottom currents than by turbidity currents.

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Rhythmic layers

FIGURE 4.27 Core photograph showing rhythmic layers of sand and mud. Middle Pleistocene, Gulf of Mexico.

Internal truncation surface

FIGURE 4.28 Core photograph showing cross-laminated sand unit (dipping to the left) with internal truncation surfaces (arrows). Middle Pleistocene, Ewing Bank Block 826, Gulf of Mexico. After Shanmugam et al. (1993a), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

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Erosional truncation surface

FIGURE 4.29 Core photograph showing an erosional truncation surface (arrow) between cross-laminated sand unit and overlying mud unit. Upper Pliocene, Ewing Bank Block 826, Gulf of Mexico. After Shanmugam et al. (1993a), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

Bidirectional crosslamination

FIGURE 4.30 Core photograph showing bidirectional cross-laminae (dashed lines). Upper Pliocene, Ewing Bank Block 826, Gulf of Mexico. After Shanmugam et al. (1993a), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

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Sand

Mud

FIGURE 4.31 Core photograph showing megascopic inverse size grading (arrow). Note the gradational nature of basal contact. Middle Pleistocene, Gulf of Mexico. After Shanmugam et al. (1993a), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

0.5 mm FIGURE 4.32 Photomicrograph of a fine-grained sand layer (arrow) showing microscopic inverse size grading and sharp upper contact. Note largest quartz grain (tip of arrow) at the top of this sand layer. Middle Pleistocene, Gulf of Mexico. After Shanmugam et al. (1993a), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

Eroded top of ripple

Preserved top of ripple

FIGURE 4.33 Core photograph showing discrete sand units comprised of current ripples with variable dip directions suggesting multiple current directions. Preserved (lower arrow) and eroded (upper arrow) tops of ripples indicate variable energy conditions of the current. Ripples with curved bases probably indicate wave influence. Middle Pleistocene, Ewing Bank Block 826, Gulf of Mexico. After Shanmugam et al. (1993a), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

FIGURE 4.34 Core photograph showing ripple cross-lamination (arrow). Middle Pleistocene, Gulf of Mexico.

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Flaser bedding (core and outcrop) (Figure 4.35). Mud-offshoots (core and outcrop) (Figures 4.36 and 4.37). Erosional basal contact (core and outcrop) (Figure 4.38). Well-sorted sand and little depositional matrix (core and outcrop).

Most of the features listed above are interpreted as the products of deposition by traction or combined traction and suspension (Figure 4.39). As with deposits of contour-following thermohaline currents, it is impossible to establish that a given sedimentary structure in the rock record was originated by wind-driven bottom currents, without establishing the paleo-water circulation pattern independently. Again, the general term “bottom-current reworked sands” is appropriate in many cases. Sand layers with traction structures occur in discrete units, but not as part of a vertical sequence of structures. Because traction structures are also observed in deposits of thermohaline-induced bottom currents (see Section 4.5.3), caution must be exercised in classifying a deposit as a “contourite” based solely on traction structures without independent evidence for contour-following bottom currents in the area.

FIGURE 4.35 Core photograph showing flaser bedding. Note the presence of mud in ripple troughs (arrow). Upper Pliocene, Ewing Bank Block 826, Gulf of Mexico. After Shanmugam et al. (1993a, reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

Mud-offshoot

FIGURE 4.36 Core photograph showing discrete sand units with current ripples and mudoffshoots (arrow). Note sigmoidal configuration of ripples and truncated tops. Middle Pleistocene, Ewing Bank Block 826, Gulf of Mexico. After Shanmugam et al. (1993a), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

FIGURE 4.37 Sketch showing discrete sand units with current ripples and mud-offshoots. After Shanmugam et al. (1995c), with permission from SEPM.

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FIGURE 4.38 Core photograph showing erosional basal contact (arrow). Middle Pleistocene, Ewing Bank Block 826, Gulf of Mexico.

4.7 DEEPWATER TIDAL BOTTOM CURRENTS I maintain a distinction between “deep-water tidal bottom currents” as a general category and “baroclinic currents” as a special category. This is necessary because not all deep-water tidal currents are baroclinic in origin. In addition, our understanding of baroclinic currents associated with internal tides is still in its infancy. Controversies and complexities associated with internal tides and ocean mixing have been addressed by Garrett (2003).

4.7.1 Submarine Canyons Deep-marine tidal bottom currents in submarine canyons and in their vicinity are one of the best-studied and most extensively documented modern geologic processes (e.g., Shepard et al., 1969, 1979; Shepard, 1976; Beaulieu and Baldwin, 1998; Petruncio et al., 1998; Xu et al., 2002). During the past four decades, an understanding of deep-marine tidal bottom currents has been achieved by synthesizing a great wealth of information that includes direct observations from deep-diving vehicles, dredging from canyon floors, underwater photographs of canyon floors, photographs and X-radiographs of box cores, seismic profiles, and current-velocity measurements (Shepard and Dill, 1966; Shepard et al., 1969, 1979; Dill et al., 1975; Shepard, 1976).

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Climbing-ripple cross-bedding

Mud-offshoots Suspension (Mud-offshoot)

Traction and suspension 5 cm

5 cm

Traction

Lenticular bedding

Flaser bedding

Suspension Traction

5 cm

5 cm

Suspension Traction

Horizontal bedding

Rhythemic bedding Sharp upper contact

Traction

5 cm

1 cm

Traction Suspension

Cross-bedding

Sharp upper contact Sharp upper contact Inverse grading

Erosion

10 cm

5 cm

Traction

Fine sand

Gradational lower contact

Mud

FIGURE 4.39 Summary of traction features interpreted as indicative of deep-water bottom current reworking. After Shanmugam et al. (1993b), Geological Society of America.

Selected examples of studies that dealt with tidal processes and/or their deposits in modern and ancient deep-water environments have been reviewed by Shanmugam (2003, 2008c). In order to discuss deep-marine tidal bottom currents, one must first establish a clear framework on submarine canyons. This is necessary because tidal currents tend to focus their energy within submarine canyons. Harris and Whiteway (2011), based on ETOPO1 (Appendix A) bathymetric grid,

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have compiled the first inventory of 5,849 separate large submarine canyons in the world’s oceans. They have classified canyons into three types, namely Type 1 shelf-incising canyons having heads with a clear bathymetric connection to a major river system (Figures 4.40 and 4.41), Type 2 shelf-incising canyons with no clear bathymetric connection to a major river system (Figures 4.42 and 4.43), and Type 3 blind canyons that incise onto the continental slope (Figures 4.42 and 4.43). In order to differentiate the role of tidal currents in different types of submarine canyons, I have further subdivided the Type 1 into Type 1A and Type 1B based on the position of canyon head. As a result, there are four types of canyons (Figure 4.44, see color plate). Their modified definitions are as follows: 1. Type 1A: Land-incising canyons having heads that incise directly into estuaries or rivers with tidal influence. In the Type 1A Zaire (formerly the Congo) Canyon in West Africa (Figure 4.45), for example, the canyon head can be traced 25 km up the estuary on land (Heezen et al., 1964; Shepard and Emery, 1973; Droz et al., 1996). The Zaire Canyon is simply a deep-water extension of the Zaire Estuary. The width and the relief of the canyon increase seaward from the estuary reaching a maximum width of 15 km and a maximum relief of 1,300 m near the shelf break (Babonneau et al., 2002). The mean tidal range in the Zaire Canyon is 1.3 m (Shepard et al., 1979). In this estuary-canyon physiographic continuum, tidal currents influence the deposition from shallowwater estuary into deep-water canyon settings. As a consequence, there are no differences between shallow-water estuarine facies (Shanmugam et al., 2000) and deep-water canyon facies (Shanmugam, 2003). This uniformity in deposition has seldom been addressed in the literature, and is of the essence in this chapter. 2. Type 1B: Shelf-incising canyons having heads with a connection to a major river or estuarine system, but they do not incise onto the land. Astoria Canyon is an example of the Type 1B with its head having a sediment-input link with the Columbia River (Figure 4.46A). The distance between the Astoria Canyon head on the shelf (130 m or 426 ft) and the mouth of the Columbia River is 17 km (11 mi). Maximum tidal range of 3.5 m has been documented at the estuarine river mouth (Hamilton, 1984). Astoria Canyon with its sinuous geometry (Figure 4.46B) has strong tidal currents that transport sediment up-canyon (Bosley et al., 2004). On the U.S. Atlantic continental slope, the Wilmington Canyon may be qualified to be a Type 1B (see “Submarine Canyon” in Appendix A). This is because the Chesapeake Bay, which lies directly upslope of the canyon head, is fed by the Susquehanna River. Keller and Shepard (1978, their Table 2-1) made direct measurements of tidal currents and reported

50

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Bottom-Current Reworked Sands

0m

10

1000

2000

3000

Zaire Canyon, West Africa

0

3,000 km

Land 100 m isobaths 1000 m isobaths

Type 1-shelf-incising, river-associated Type 2-shelf-incising Type 3-blind (headless)

FIGURE 4.40 Type 1 shelf-incising, river-associated Zaire (formerly Congo) Canyon. Note that I have reclassified this canyon as Type 1A in this book. It is isolated with the nearest adjacent canyons about 100 km away. Compiled from Harris and Whiteway (2011), with permission from Elsevier.

169

170

0

200 km

100 m 1000

2000

3,000 km

Land 100 m isobaths 1000 m isobaths

Type 1-shelf-incising, river-associated Type 2-shelf-incising Type 3-blind (headless)

FIGURE 4.41 Type 1 shelf-incising, river-associated Swatch-No-Ground Canyon adjacent to the Ganges-Brahmaputra River, Bay of Bengal. Note that I would reclassify this canyon as Type 1B. It is also isolated with nearest canyon about 100 km away. Compiled from Harris and Whiteway (2011), with permission from Elsevier.

New Perspectives on Deep-water Sandstones

Swatch-No-Ground Canyon, Bay of Bengal

0

Chapter | 4

0

50 100 km

2000

1000

0

3,000 km

Bottom-Current Reworked Sands

100 m

Canyons on the slope of Gulf of Lion

Land 100 m isobaths 1000 m isobaths

Type 1-shelf-incising, river-associated Type 2-shelf-incising Type 3-blind (headless)

FIGURE 4.42 Types 2 and 3 canyons on the slope of the Gulf of Lion, northern Mediterranean Sea. They are spaced less than 10 km apart from each other. Compiled from Harris and Whiteway (2011), with permission from Elsevier.

171

172

200 km

0

La

ur

en

tia

nc

100

ha

nn

m

el

2000

100

0

0m

0

3,000 km

Canyons on the Canadian continental margin

Land 100 m isobaths 1000 m isobaths

Type 1-shelf-incising, river-associated Type 2-shelf-incising Type 3-blind (headless)

FIGURE 4.43 Types 2 and 3 canyons near the Laurentian Channel, many of which incise the shelf, incised into the glacial trough mouth fan. Compiled from Harris and Whiteway (2011), with permission from Elsevier.

New Perspectives on Deep-water Sandstones

40

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173

Bottom-Current Reworked Sands

Land Estuary

Shelf

Slope

Basin

Surface Internal tide tide Common Rare

Type 1A

River Type 1B

Type 2

Type 3

Rare Common (A)

(B)

(C)

FIGURE 4.44 (A) Proposed classification of canyons for understanding tidal energy. This is a modified scheme of Harris and Whiteway (2011). Type1A land-incising canyon into estuary or river system with tidal influence. Type 1B shelf-incising canyons having heads with connection to a major river or estuarine system, but they do not incise onto the land. Type 2 shelf-incising canyons with no clear connection to a major river or estuarine system. Type 3 slope-incising blind canyons with their heads confined to the continental slope. (B) Increasing influence of surface (barotropic) tides from Type 3 to Type 1A canyons. (C) Increasing influence of internal (baroclinic) tides from Type 1A to Type 3 canyons.

FIGURE 4.45 Type 1A Zaire (formerly Congo) submarine canyon. Note the canyon head can be traced 25 km up the estuary on land. After Heezen et al. (1964), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

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New Perspectives on Deep-water Sandstones

Canyon head

Astoria Canyon

49°N VI

(B) 48°N WA 47°N Astoria canyon 46°N

45°N

Columbia river OR

44°N 127°W 126°W 125°W 124°W 123°W

(A)

FIGURE 4.46 (A) Map showing bathymetry and earthquakes (white dots) located near Astoria Canyon and the surrounding continental margin. Yellow, ,1,000 m deep; green, 1,000 2,000 m deep; blue, 2,000 m deep. The 79 earthquakes shown in this map occurred between August 1991 and January 2001. VI, Vancouver Island, WA, Washington, OR, Oregon. Image courtesy of Lewis and Clark 2001, NOAA/OER. Astoria Canyon: A Natural Laboratory, Bob Embley, Geophysicist, http://oceanexplorer.noaa.gov/explorations/lewis_clark01/background/geology/ media/bathymetry.html. Accessed July 10, 2011. (B) Type 1B Astoria Canyon (looking southeast) showing sinuous geometry. ROPOS dive tracks are superimposed: blue, June 28; fuschia, June 29; green, June 30; white, July 1; red and black, July 2; and yellow, July 3. Relief is 400 500 m from the top of the wall to the canyon bottom. The entire canyon is approximately 120 km long “as the crow flies” and has an approximate area of 1,000 km2 (using 8 km as the average width). Image courtesy of Lewis and Clark 2001, http://oceanexplorer.noaa.gov/explorations/lewis_clark01/logs/jul03_scisum/media/threedastoria.html. Accessed July 10, 2011.

maximum velocities of 44 cm s21 at a depth of 723 m in the Wilmington Canyon. During the 19-day survey period, the net current flow was upcanyon, confirming the tidal origin of these currents. 3. Type 2: Shelf-incising canyons with no clear connection to a major river or estuarine system. Many canyons on the slope of the Gulf of Lion, northern Mediterranean Sea are of this type. Canyon heads are commonly located near the shelf edge, nearly 75 km away from the shoreline (Figure 4.42). 4. Type 3: Slope-incising blind canyons with their heads confined to the continental slope. Many canyons associated with the Laurentian Channel on the Canadian Atlantic margin are of this type. Canyon heads are

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Bottom-Current Reworked Sands

175

located invariably on the slope, nearly 400 km away from the shoreline (Figure 4.43). Although the position of canyon heads near the shelf edge or on the upper slope in Types 2 and 3 tends to reduce the chances of strong tidal currents entering the canyons from shallow-water estuaries, there are examples that Type 2 and 3 canyons can serve as favorable settings for barotropic tidal currents. For example, in the present-day Krishna Godavari (KG) Basin, Bay of Bengal, tidal estuaries, narrow shelf widths (15 25 km), and upper-slope canyons are considered to be viable factors for focusing tidal currents in submarine canyons during the Pliocene deposition of reservoir sands (Shanmugam et al., 2009). Another example is Hervey Bay and Fraser Island, southeast Australia, where transport of coastal sand into the deep ocean by ebb tidal currents during the present highstand has been documented (Boyd et al., 2008). The Hervey Bay area in Australia is characterized by (a) mesotidal range (2 4 m) (Boyd and Leckie, 2004), (b) high-velocity (150 cm s21) ebb tidal currents, (c) northward-flowing longshore currents, (d) narrow (5 25 km) shelf widths, and (e) numerous slope gullies of Type 3 canyons. In both the KG Basin and the Hervey Bay examples, narrow shelf is the key factor. So far, I have discussed only tidal currents induced by surface (barotropic) tides in submarine canyons. These currents are more common in Type 1A than in deep-water Type 3 canyons (Figure 4.44, see color plate). On the other hand, baroclinic currents induced by internal tides, discussed in Section 4.8 and by Allen and Durrieu de Madron (2009), are expected to be more common in deep-water Type 3 canyons than in Type 1A canyons (Figure 4.44, see color plate).

4.7.2 Current Velocity Shepard et al. (1979) measured current velocities in 25 submarine canyons worldwide at water depths ranging from 46 to 4,200 m by suspending current meters, usually 3 m above the sea bottom (Figure 4.47). Shepard et al. (1979) also documented systematically that up- and down-canyon currents closely correlated with timing of tides (Figure 4.48). These canyons include the Hydrographer, Hudson, Wilmington, and Zaire in the Atlantic Ocean; and the Monterey, Hueneme (Figure 4.48), Redondo, La Jolla/Scripps, and Hawaii canyons in the Pacific Ocean. Maximum velocities of up- and downcanyon currents commonly ranged from 25 to 50 cm s 1 (Table 4.3). Keller and Shepard (1978) reported velocities as high as 70 75 cm s21, velocities sufficient to transport even coarse-grained sand, from the Hydrographer Canyon. In the Niger Delta area of West Africa, five modern submarine canyons (Avon, Mahin, Niger, Qua Iboe, and Calabar) have been recognized (Figure 4.49). In the Calabar River, the tidal range is 2.8 m and tidal flows

Submarine Canyon Depth: 46 – 4,200 m

Sea level

Ebb

Flood

Current meter 3m above base

Velocity range: 25 – 50 cm s–1

FIGURE 4.47 Conceptual diagram showing cross-section of a submarine canyon with ebb and flood tidal currents (opposing arrows). Shepard et al. (1979) measured current velocities in 25 submarine canyons at water depths ranging from 46 to 4,200 m by suspending current meters commonly 3 m above the sea bottom. Measured maximum velocities commonly range from 25 to 50 cm s 1. After Shanmugam (2003), with permission from Elsevier.

30

30 cm s–1

20

20 cm s–1

10

10 cm s–1

cm s–1

Downcanyon

Upcanyon

Hueneme Sta. 28 Lo. 1 Depth 448 m 3 mAB V vs Time

0

Tide

10

10 cm s–1

20

20 cm s–1

30

30 cm s–1

2m

1

0 2000 hr. 12 2/12/73

24

36

48 Hours

60

72

84

FIGURE 4.48 Time/velocity plot of data obtained at 448 min the Hueneme Canyon, California, showing excellent correlation between the timing of up- and down-canyon currents and the timing of tides obtained from tide tables (solid curve). 3 mAB=velocity measurements were made 3 m above sea bottom. Modified after Shepard et al. (1979), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

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TABLE 4.3 Selected Examples of Maximum Velocities of Up- and DownCanyon Currents Measured at Varying Water Depths by Suspending Current Meters 3 m Above Sea Bottom Downcanyon Current Velocity (cm s21)

Water Depth (m)

Submarine Canyon and Location

Mean Tidal Range (m)

Up-canyon Current Velocity (cm s21)

348

Hydrographer, Atlantic, U.S.

1.6

39

52

375

La Jolla, Pacific, U.S.

2.5

19

18

400

Zaire (Congo), West Africa

1.3

22

13

448

Hueneme, Pacific, U. S.

2.4

32

32

458

Santa Monica, Pacific, U.S.

2.7

27

30

914

Wilmington, Atlantic, U.S.

1.8

20

21

1445

Monterey, Pacific, U. S.

2.0

30

30

1737

Kaulakahi, Pacific Islands

0.9

26

24

1904

Rio Balsas, Mexico

0.7

21

21

Source: Compiled from Shepard et al. (1979).

with reversible currents are common (Allen, 1965). In the Calabar Estuary, maximum ebb-current velocities range from 60 to 280 cm s 1, and flood current velocities range from 30 to 150 cm s 1. These velocities are strong enough to transport particles of sand and gravel size. The Calabar Estuary has a deep-water counterpart that cuts through sediments of the outer shelf and slope, forming the modern Calabar Submarine canyon (Figure 4.49). Thus, as they do in the Zaire Canyon to the south, tidal currents are likely to operate in the Calabar and Qua Iboe Canyons.

4.7.3 Sedimentological Criteria Sedimentary features indicative of tidal processes in shallow-water environments have been well established (e.g., Reineck and Wunderlich, 1968;

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New Perspectives on Deep-water Sandstones

FIGURE 4.49 Location of modern and ancient submarine canyons in the Gulf of Guinea, West Africa. Outside submarine canyons, the shelf-slope break (dashed line) is not only an important physiographic boundary between shelf and slope but also a major controlling factor of processes on the shelf (e.g., tides and waves) and on the slope (e.g., mass transport). However, within submarine canyons (e.g., recent Calabar Canyon), the shelf-slope break does not control processes. Map modified after Petters (1984), reproduced with permission from Blackwell Publishing.

Klein, 1970; Visser, 1980; Terwindt, 1981; Allen, 1982; Banerjee, 1989; Nio and Yang, 1991; Dalrymple, 1992; Archer, 1998; Shanmugam et al., 2000). Traction structures that develop in shallow-water estuaries also develop in deep-water canyons and channels with tidal currents (Shanmugam, 2003). Examples used in this section are composed of conventional core from offshore Nigeria (Figure 4.50), Equatorial Guinea (Figure 4.51), Bay of Bengal (Chapter 7), and outcrop from SE France (Figure 4.52). General characteristics of deep-water tidal deposits are as follows: G G

G

G

Heterolithic facies (core and outcrop). Rhythmic alternation of sand/mud couplets (tidal rhythmites) (core and outcrop) (Figures 4.53 and 4.54A). Double mud layers (core and outcrop) (Figures 4.54B, 4.55, 4.56, and 4.57): These structures are unique to deposition from tidal currents (Visser, 1980). Thick thin sand layers with double mud layers (core and outcrop) (Figure 4.58). Visser (1980) originally explained the origin of double mud layers by alternating ebb and flood tidal currents with extreme timevelocity asymmetry in shallow-water subtidal settings. The thick sand units likely reflect deposition during dominant tides, whereas the thin sand units are probably products of subordinate tides. The mud layers

Chapter | 4 Bottom-Current Reworked Sands

179

FIGURE 4.50 Location map of Edop Field, offshore Nigeria. From Shanmugam (1997b).

180

New Perspectives on Deep-water Sandstones

FIGURE 4.51 Location map showing the Zafiro Field, Equatorial Guinea. Cored wells: Z 2, Zafiro 2; Z 3, Zafiro 3, Z 4, Zafiro 4. After Famakinwa et al., 1996; see also Shanmugam et al., 1997b.

FIGURE 4.52 Map showing location of Annot Sandstone (Eocene-Oligocene) units (solid dots) measured along a road section near Peira Cava, French Maritime Alps, north of Nice, southeastern France. Road section is partly based on Lanteaume et al. (1967). After Shanmugam (2002a), with permission from Elsevier.

Chapter | 4

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Bottom-Current Reworked Sands

FIGURE 4.53 Core photograph showing rhythmic bedding of rippled fine-grained sand and mudstone. Zafiro Field, Pliocene, Equatorial Guinea.

cm. Sand

DML DML

Mud Contorted layer DML

DML

Mudstone (massive) clasts

DML

Silty layer

Ripples

Rhythmites

DML

Mud Draped Ripples Double mud layers (DML)

DML DML

Double Mud Layers (DML)

DML

(A)

DML

Burrow C

Carbonaceous debris Burrows

(B)

Sandstone fragments

FIGURE 4.54 (A) Core photograph showing rhythmic bedding (rhythmites) in mudstone with double mud layers (DML). Silty layers are light gray. (B) Core photograph showing DML in mudstone. Silty layers are light gray. After Shanmugam et al. (2009), with permission from SEPM.

G

represent deposition of mud from settling during slack-water periods that occur during intervening periods between depositions of rippled sand from traction (Figure 4.59). Thick (spring)/thin (neap) bundles (core and outcrop) (Figure 4.56).

DML

FIGURE 4.55 Core photograph showing double mud layers (arrow) in Pliocene sand. Edop Field. After Shanmugam (2003), with permission from Elsevier.

20% 60% 80%

m C

2072

S VFS FS

MS Double mud layers (DML)

Well 2

DML

2073

Core 8

N

2074 S DML in mudstone

2075

N

Increasing mud

2076

Sandstone fragment

2077

Mudstone clasts

(A)

(B)

FIGURE 4.56 (A) Sedimentological log showing alternation of sand and mudstone intervals with continuous presence of double mud layers (DML). Note floating sandstone rock fragments and mudstone clasts in a basal mudstone interval. (B) Core photograph showing rhythmic bedding (rhythmites) and double mud layers (DML, arrows) in sand. N, Neap (thin) bundle; S, Spring (thick) bundle. After Shanmugam et al. (2009), with permission from SEPM.

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Bottom-Current Reworked Sands

DML

FIGURE 4.57 Outcrop photograph of Annot Sandstone (Eocene-Oligocene) showing double mud layers (DML, arrows). Note mud-draped ripples between upper and middle arrows. After Shanmugam (2003), with permission from Elsevier.

DML

Thick Thin

Thick-thin sand layers

FIGURE 4.58 Core photograph showing thick thin sand layers with double mud layers (DML) in Pliocene sand. Edop Field, Pliocene, offshore Nigeria.

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New Perspectives on Deep-water Sandstones

(A) Dominant current stage (Traction)

(B) First slack-water stage (Settling) Settling

(C) Subordinate current stage (Traction)

(D) Second slack-water stage (Settling)

Double mud layers FIGURE 4.59 Schematic diagram showing origin of double mud layers. Original diagram by Visser 1980, courtesy of Geological Society of America. Diagram slightly modified using the style of Reading and Collinson (1996). G

G G

Climbing ripples; cross-beds with mud-draped foresets (core and outcrop). Bidirectional (herringbone) cross-bedding (core and outcrop). Sigmoidal cross-bedding (i.e., full-vortex structures) with mud drapes and tangential basal contacts (core and outcrop) (Figures 4.60 and 4.61). These sigmoidal cross-beddings of the Annot Sandstone from SE France were not recognized by Bouma (1962) who originally interpreted these features as turbidites. I have reinterpreted these features as tidal bundles of deep-marine tidal currents. In sigmoidal cross-bedding, tangential basal contacts, steeply dipping foresets, and fanning of the foresets may be equivalent to the full-vortex part of tidal bundles in shallow-water environments (Terwindt, 1981). Tidal bundles represent a lateral succession of cross-strata deposited in one event by the dominant tide (Terwindt, 1981). During Annot deposition, tidal currents were probably responsible for forming sigmoidal cross-bedding in canyon or channel settings. Tangential toesets are not unique to tidal deposits because these structures only indicate that the avalanche facies resulted from a combination of tractional and suspension fall-out processes. Such processes are common in other environments as well (e.g., fluvial and eolian). The deep-water

Chapter | 4 Bottom-Current Reworked Sands

FIGURE 4.60 (A) Sedimentological log of Annot Sandstone (Unit 11) showing sigmoidal cross-bedding with mud drapes. Note underlying inverse grading of gravel layer. Note mud-draped ripple beds in adjacent mudstone intervals. (B) Outcrop photograph of Annot Sandstone (Unit 11) showing sigmoidal cross-bedding with tangential basal contacts. Upper arrow shows stratigraphic position of sigmoidal cross-bedding and lower arrow shows stratigraphic position of inverse grading. See Figure 4.52 for location of Unit 11. After Shanmugam (2003), with permission from Elsevier.

185

186 New Perspectives on Deep-water Sandstones

FIGURE 4.61 (A) Sedimentological log of Annot Sandstone (Unit 10) showing three sets of sigmoidal cross-bedding. Note contorted layers of a sandstone bed above (4.75 m) and a normally graded sandstone bed with clasts below (1 2 m). Note rippled sand beds in intervening mudstone intervals. (B) Outcrop photograph of Annot Sandstone (Unit 10) showing the middle set of sigmoidal cross-bedding (arrow) with mud drapes (dark coloration). Note faint tangential lower contacts of the upper set. Lower set of sigmoidal cross-bedding is poorly developed and thus is not obvious in the photograph. Rarely, mica has been observed along with mud. Arrow shows stratigraphic position of photograph. See Figure 4.52 for location of Unit 10. After Shanmugam (2003), with permission from Elsevier.

Chapter | 4

Bottom-Current Reworked Sands

187

FIGURE 4.62 A model for tidal bundles in shallow-marine environments. This model has been adopted to explain sigmoidal cross-bedding in deep-water Annot Sandstone in SE France in this book. Original diagram from Terwindt (1981). Modified by Banerjee (1989), and further simplified by Shanmugam et al. (2000).

G G G G G G

origin of the Annot Sandstone has restricted our options to either turbidity currents or tidal currents. Under this constraint, tangential toesets can be explained better by tidal currents than by turbidity currents. This is because a genetic link between tangential cross-bedding and tidal currents has been well established (Terwindt, 1981) (Figure 4.62), whereas a link between tangential cross-bedding and turbidity currents has not been documented. Lee et al. (2002) reported the generation of upslope-migrating antidunes by Froude-supercritical turbidity currents. However, the origin of sigmoidal cross-bedding with mud drapes is difficult to explain by upslope-migrating antidune bedforms deposited by turbidity currents. Furthermore, no one has ever generated sigmoidal cross-bedding with mud drapes by turbidity currents in flume experiments. Nor has anyone documented sandy turbidity currents in modern oceans (Section 1.2.1). Other outcrop examples with deep-water sigmoidal cross-beddings of tidal origin have been reported by Mutti et al. (1984b) and Neumeier (1998). Alternation of parallel and cross-laminated units (core and outcrop). Reactivation surfaces (core and outcrop). Crinkled laminae (core and outcrop). Elongate mudstone clasts (core and outcrop). Flaser bedding (core and outcrop) (Figure 4.63). Wavy bedding (core and outcrop).

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New Perspectives on Deep-water Sandstones

Flaser bedding

FIGURE 4.63 Core photograph showing flaser bedding in Pliocene sand. Edop Field, offshore Nigeria. G G G G

G

G

Lenticular bedding (core and outcrop) (Figure 4.64) Mud-draped ripples (Figure 4.65) Current ripples (core and outcrop) (Figure 4.66). Wave ripples (core and outcrop) (Figure 4.67). This structure has conventionally been considered to form only in shallow-water environments, but in this case it occurs in deep-water lithofacies. Horizontal lamination with gradational upper contact (core and outcrop) (Figure 4.68). Alternating traction and suspension structures (core and outcrop). Klein (1975), based on studies of DSDP (Leg 30, Sites 288 and 289) cores, suggested that current ripples, micro-cross-laminae, mud drapes, flaser bedding, lenticular bedding, and parallel laminae reflect alternate traction and suspension deposition from tidal bottom currents in deep-marine environments. In short, deep-water tidalites have been documented both in the modern settings and in the ancient rock record unequivocally.

4.7.4 Facies Associations in Submarine Canyons Submarine canyons are not only unique for providing a protected environment for focusing tidal energy from shallow-marine estuaries to deep-marine canyons but also prone to generating mass movements (e.g., slides, slumps, and debris flows) due to failure of steep canyon walls. Recognition of tidal facies

Chapter | 4

189

Bottom-Current Reworked Sands

Top

Lenticular bedding and starved ripples

Bottom

FIGURE 4.64 Core photograph showing lenticular bedding and starved ripple (arrow) in Pliocene sand. Edop Field, offshore Nigeria.

Mud-draped ripples

FIGURE 4.65 Core photograph showing mud-draped ripples in Pliocene sand. Edop Field, offshore Nigeria.

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New Perspectives on Deep-water Sandstones

Current ripples

FIGURE 4.66 Core photograph showing current ripples in Pliocene sand. Edop Field, offshore Nigeria.

Wave ripple

FIGURE 4.67 Core photograph showing wave ripple in Pliocene sand. Edop Field, offshore Nigeria.

Chapter | 4

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Bottom-Current Reworked Sands

in deep-water sequences is important in understanding sand distribution because deposits of tidal processes and mass transport processes (i.e., slides, slumps, and debris flows) characterize fills of modern and ancient submarine canyons (Table 4.4). This complex facies association (Figure 4.69), mimicking both shallow-water and deep-water deposits, has been recognized in the modern La Jolla canyon box cores (offshore California), ancient Qua Iboe

Horizontal laminae

FIGURE 4.68 Core photograph showing horizontal lamination with gradational upper contact in Pliocene sand. Edop Field, offshore Nigeria. Note faint normal grading.

TABLE 4.4 Summary of Examples Showing a Close Association of Tidal Currents and Mass Flows and Their Deposits in Modern and Ancient Submarine Canyons Modern and Ancient Submarine Canyons

Tidal Currents and Their Deposits

Mass Flowsa and Their Deposits

Modern, La Jolla, Pacific, U.S.

Current reversal (Shepard et al., Slide and debris flow 1979; their Table 1) and (Shepard et al., 1969, their double mud layers (Shepard Figures 9, 10, and 21B) et al., 1969, their Figure 19 A; this study)

Modern, Scripps, Pacific, U.S.

Current reversal (Shepard et al., Slumps (Marshall, 1978) 1979, their Table 1) (Continued )

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New Perspectives on Deep-water Sandstones

TABLE 4.4 (Continued) Modern and Ancient Submarine Canyons

Tidal Currents and Their Deposits

Mass Flowsa and Their Deposits

Modern, Monterey, Pacific, U.S.

Current reversal (Shepard et al., Mass movements (Martin 1979, their Table 1) and Emery, 1967, p. 2281)

Modern, Pribilof, Bering Sea

Tidal currents (Sapozhnikov et al., 2011)

Mass movements (Carlson and Karl, 1984)

Modern, Navarinsky, Bering Sea

Internal waves and tides (Karl et al., 1986)

Mass movements (Carlson et al., 1982)

Modern, Rio Balsas, Pacific, Mexico

Current reversal (Shepard et al., Slumps (Dill et al. 1975) 1979, their Table 1)

Modern, San Lucas, Pacific, Baja Mexico

Current reversal (Shepard et al., Grain flow and sandy debris 1979, their Table 1) and flow (Shepard and Dill, Interference ripples (Shepard 1966, their Figures 55 &139) and Dill, 1966, their Figure 53C)

Modern, Hudson, Atlantic, U.S.

Current reversal (Shepard et al., Muddy debris flow (Shepard 1979, their Table 2) and Dill, 1966, their Figure 74)

Modern, Zaire (Congo), Atlantic, West Africa

Current reversal (Shepard et al., Slumping (Shepard and 1979, their Table 1) Emery, 1973, their Figure 7)

Modern, “Swatch of No Ground,” Bay of Bengal

Tidal currents (Michels et al., 2003)

Slides and slumps (Michels et al., 2003)

Modern, Mera, Honshu, Japan

Tidal currents and sigmoid ripples (Shepard and Dill, 1966; see also Tsuji, 1993 for tidal currents with 51 cm s21 velocity in the Ryukyu Trench to the south)

Breccia recovered from dredging (Shepard and Dill, 1966), which could be interpreted as deposits of mass flows

Pliocene, Qua Iboe, Atlantic, West Africa

Double mud layers (this study)

Slumping and debris flow (this study)

Pliocene, Krishna Godavari Basin, Bay of Bengal

Double mud layers (Shanmugam et al., 2009)

Sandy debris flows (Shanmugam et al., 2009)

Eocene-Oligocene, Annot Sandstone, SE France (Outcrop study)

Double mud layers and sigmoid cross-bedding (this study)

Slumping and debris flow (Stanley et al., 1978; his Figure 8-3, this study)

(Continued )

Chapter | 4

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Bottom-Current Reworked Sands

TABLE 4.4 (Continued) Modern and Ancient Submarine Canyons Eocene, Torrey, California (Outcrop study)

Tidal Currents and Their Deposits Rippled and cross-bedded facies (May et al., 1983)

Mass Flowsa and Their Deposits Slides, sandy debris flows, grain flows (May et al., 1983)

Source: Updated after Shanmugam (2003), with permission from Elsevier a Mass flows include slides, slumps, grain flows, and debris flows, but not turbidity currents (see Dott, 1963). Note: This table is based on published information on direct observations from deep-diving vehicles, underwater photographs of canyon floors, photographs and x-radiographs of box cores, current-meter data, seismic profiles, and study of conventional core and outcrop data. Note that two of the published cores (La Jolla) and outcrops (Annot) have been reinterpreted in this study.

Canyon-fill facies Double mud layers (Tidal) Floating mudstone clasts and quartz granules (Debris flows)

Contorted layers and floating quartz granules (Slumps)

Double mud layers (Tidal) Mud-draped ripples (Tidal)

FIGURE 4.69 Facies association showing interbedded occurrence of double mud layers (tidal origin), floating mudstone clasts and quartz granules (debris-flow origin), double mud layers (tidal origin), contorted layers and floating quartz granules (slump origin), and double mud layers with mud-draped ripples (tidal origin). In the rock record, such a facies association may be used as evidence for deposition within submarine canyons. Modified after Shanmugam (2003), with permission from Elsevier.

Canyon conventional cores (Pliocene, Edop Field, offshore Nigeria), ancient Pliocene canyon conventional cores (KG Basin, Bay of Bengal), and ancient Annot Sandstone outcrops (Eocene Oligocene, onshore SE France), among others. It appears that the association of tidal and mass flow facies is unique to

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canyon environments. Therefore, this facies association may be used as a criterion for inferring submarine canyon settings in the rock record where direct evidence for canyon filling is lacking. Because MTD (i.e., deposits of slides, slumps, and debris flows) can occur both inside and outside submarine canyons, the correct identification of tidal facies in deep-water sequences is extremely critical in establishing the facies association.

4.8 BAROCLINIC CURRENTS (INTERNAL TIDES) 4.8.1 Nomenclature and Background A plethora of nomenclature is in use for internal waves and tides. In particular, the term baroclinic has been used by different authors with different meanings (Wunsch, 1996). Thus, it is useful to explain these terms to minimize confusion. Surface waves, caused by wind (meteorological force) blowing over the water surface (Komar, 1976), develop at the interface between water and air. Internal waves (Apel, 2002), recognized since the Viking times (Ekman 1904), are gravity waves that oscillate along the interface between two water layers of different densities (Figure 4.70). These waves are common phenomena in coastal seas, fjords, lakes, and the atmosphere. In shallow-water shelf environments, waters can range from well mixed to density-stratified types. Most shelf waters are vertically well mixed. In deep-water environments, on the other hand, most of the ocean is vertically stratified, with an upper lowdensity layer and a lower high-density layer (Figure 4.70). The interface between layers of different densities (i.e., pycnocline) can be caused either by differences in temperature (i.e., thermocline) or by salinity (i.e., halocline).

Shallow water

Deep water

Surface waves (meteorological)

Sea level

Internal waves & tides (astronomical)

Warm layer

Shelf

Thermocline Cold layer Basin

FIGURE 4.70 Schematic diagram showing meteorological surface waves in shallow-water shelf environments and astronomical internal waves and tides in deep-water slope and basinal environments. Internal waves and tides, occurring along a pycnocline (thermocline), can extend onto shelf environments. Internal waves and tides typically have much higher amplitudes than surface waves. Diagram based on concepts of Inman et al. (1976) and Maxworthy (1979). Not to scale. Modified after Shanmugam (2008b), with permission from Elsevier.

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Navrotsky et al. (2004) made observations of internal waves and spatial inhomogeneities of a thermohaline structure during the warm season on the continental slope and shelf of the Sea of Japan. Santek and Winguth (2005) documented internal waves, induced by the 2004 Indian Ocean tsunami, along the eastern continental slope off Sri Lanka. NASA has documented both surface waves and internal waves on the Indian Ocean waters near the Andaman Islands in a satellite image (Figure 4.71). In this image, smaller surface waves propagate in a north south direction, whereas larger internal waves with a wavelength of 5 km propagate in an east west direction, at nearly right angle to surfaces waves (Figure 4.71). Internal waves typically have much lower frequencies and higher amplitudes than surface waves because the density difference between two water layers is typically much less than the density difference between water and air. Unlike surface waves, internal waves can stretch tens of kilometers in length (Figure 4.72). They can move throughout the ocean for several hours. They have their greatest wave height at intermediate depths and

Andaman sea Surface

waves

es

av

lw

na

er

Int

Barren Island Volcano

5 km

FIGURE 4.71 Satellite image showing both surface waves (near horizontal faint waves, left) and internal waves (long diagonal waves, right) on the Indian Ocean near Andaman Islands. Large internal waves have a wavelength of 5 km. Note the active Barren Island Volcano emitting steam on the lower left of the image. This Advanced Land Imager (ALI) on the Earth Observing 1 satellite acquired the image on March 6, 2007. NASA Earth Observatory image created by Jesse Allen and Robert Simmon, using EO-1 ALI data provided courtesy of the NASA EO-1 team. Caption by Holli Riebeek. NASA Earth Observatory image credit, http://earthobservatory. nasa.gov/IOTD/view.php?id=44567. Accessed July 12, 2011.

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South China Sea Palawan

Tubbalaha Reefs

Sulu Sea

75 km Internal Waves

FIGURE 4.72 Satellite image of internal waves in the Sulu Sea between the Philippines (to the northeast) and Malaysia (to the southwest). Sunlight highlights delicate curving lines of internal waves moving to the north toward Palawan Island. The Sulu Sea is stratified with water layers of differing densities. Unlike surface waves, internal waves can stretch tens of kilometers in length and move throughout the ocean for several hours. This true-color Aqua Moderate Resolution Imaging Spectroradiometer (MODIS) image was acquired on April 8, 2003. Image courtesy Jacques Descloitres, MODIS Land Rapid Response Team at NASA/GSFC. Uniform Resource Locator (URL), http://earthobservatory.nasa.gov/Newsroom/NewImages/images.php3? img_id=15334. Accessed May 12, 2007.

their greatest velocities at the bottom (LaFond, 1962). Internal waves are of significance not only for maintaining ocean circulation by downward mixing of heat but also for sustaining biological productivity by supplying nutrients. Internal waves that correspond to periods of tides are called internal tides (Shepard, 1975). In other words, internal tides are nothing more than internal waves at a tidal frequency. According to Garrett and Kunze (2007, p. 57), “Internal tides are internal gravity waves generated in stratified waters by the interaction of barotropic tidal currents with variable bottom topography.” It is important to distinguish the surface (barotropic) tide from the

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internal (baroclinic) tide. This is because sedimentological aspects of deposits of surface tides are well established (Alexander et al., 1998), whereas those of internal tides are not. The passage of tropical cyclone Bobby in 1995 over the Western Australian Shelf influenced the generation of internal tides at 300 m depth (Davidson and Holloway, 2003). The currents associated with internal tides are termed baroclinic currents or motions. These tidal currents tend to create turbulent mixing in the stratified water column as well as at the seafloor. The term near-inertial wave has been used for internal waves that correspond to periods of upwelling (Federiuk and Allen, 1996). Internal solitary waves, consisting of a single isolated wave, are common in stratified fluids. The term soliton is commonly used as a synonym for solitary waves. Although the term soliton is strictly a mathematical solution of a nonlinear internalwave theory, it has become common practice in the offshore drilling industry to use the term to describe observations of large-amplitude nonlinear internal waves in the ocean (Hyder et al., 2005). The origin of internal solitary waves has been attributed to a number of causes, including seafloor topography (Maxworthy, 1979; Farmer and Armi, 1999). Because internal solitons are hazardous to offshore drilling operations (Fraser, 1998), a clear understanding of these internal waves has practical implications for the cost and the safety of offshore drilling operations. Gao et al. (1998) convey a simple message that all deep-water tidal currents are baroclinic in origin (i.e., internal tide related). The reality is somewhat more complex. For example, transport of coastal sand into the deep ocean by ebb tidal currents has been documented in the offshore areas of Hervey Bay and Fraser Island, southeast Australia (Boyd et al., 2008).These ebb tidal currents are not baroclinic currents; they are simply barotropic ebb tidal currents that happen to transport sand across the entire shelf and supply the sand into submarine canyons and gullies into deep water. One needs to make this distinction between barotropic ebb tidal currents, which go from the coast into the deep water, from baroclinic tidal currents that are generated in the deep water along the pycnoclines near the shelf edge (Figure 4.70). Although we know much about the surface tides and related barotropic currents, we do not know enough about the internal tides and baroclinic currents. The fate of the dissipated energy of internal tides is still unknown (Garrett and Kunze, 2007). However, both shallow-water barotropic and deep-water baroclinic currents are comparable in strength (Garrett and Kunze, 2007).

4.8.2 Submarine Canyons Allen and Durrieu de Madron (2009) observed that canyons that do not enter into large estuaries do not tend to have tidal currents along the axis of the canyon, but rather across the canyon and in the direction parallel to the shelf

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break. Several long canyons of this type have been observed to have high to very high internal tidal energy. Selected examples are as follows: G G G

G G

Hydrographer Canyon, U.S. Atlantic (Wunsch and Webb, 1979), Hudson Canyon, U.S. Atlantic (Hotchkiss and Wunsch, 1982), Monterey Canyon, U.S. Pacific (Petruncio et al., 1998; Kunze et al., 2002), La Linea Canyon, Western Alboran Sea (Lafuente et al., 1999), Gaoping Canyon, Taiwan (Lee et al., 2009).

Canyons can act to enhance tidal and internal-wave energy due to focusing (Wunsch and Webb, 1979) or due to large regions of critical slope (Hotchkiss and Wunsch, 1982). It has also been proposed (based on observations in Monterey Canyon) that small-scale roughness in the canyon is responsible for much of the internal tidal energy (Kunze et al., 2002). Shepard et al. (1979) measured velocities of up- and down-canyon currents at water depths ranging from 46 to 4,200 m. Most of these measurements were made at about 3 m above the canyon floor. Clearly, these are surface-tide induced currents. Shepard and Marshall (1978) also documented a few cases (e.g., Carmel Canyon), there is evidence for up-canyon advancing of internal waves along the canyon axis. But they were not able to demonstrate that all their current measurements were related to advancing internal waves. This raises the question whether the velocity measurements made at 4,200 m depth, close to the seafloor, were indeed representing density stratification within the canyon (i.e., pycnocline). We do not know the answer. Therefore, one should address this issue of density stratification in a given area under investigation for deep-water tidal currents. Otherwise, one will get the false impression that all deep-water tidal currents are associated with internal tides. It is important to note that not all deep-water tidal currents have been documented to be baroclinic currents using temporal offset of arrival time of internal waves in closely spaced stations. My view is that the influence of bottom-current reworking by barotropic currents may decrease from shallow-water canyons (Type 1A) to deep-water canyons (Type 3), whereas the reworking by baroclinic currents may decrease from deep-water (Type 3) canyons to shallow-water (Type 1A) canyons (Figure 4.44). However, our understanding of deposits of baroclinic currents in the rock record is meager (see discussion below).

4.8.3 Current Velocity Deep-water bottom currents have been attributed to internal waves in offshore California (Emery, 1956). Wunsch (1969) proposed amplification of

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near-bottom velocities as internal waves propagate over a shoaling bottom. Brandt et al. (2002) reported results of high-resolution velocity measurements carried out by means of a vessel-mounted acoustic Doppler current profiler on the November 12, 2000 in the equatorial Atlantic, at 44 W between 4.5 N and 6 N. The data showed the presence of three large-amplitude internal solitary waves. The pulse-like intense solitary disturbances propagated perpendicular to the Brazilian Shelf, toward the north northeast. These internal waves were characterized by maximum horizontal velocities of about 200 cm s21 and maximum vertical velocities of about 20 cm s21. Shepard et al. (1979), who studied bottom currents in submarine canyons, documented that internal waves advance in both up- and down-canyon directions. Measured values of velocity reach up to 100 cm s21 in the up-canyon direction and 265 cm s21 in the down-canyon direction. Shepard (1975) suggested that internal waves, which occur in canyon depths of up to 3,500 m, were mostly tidal in origin (i.e., internal tides). In a stratified ocean, internal tides are generated commonly above an area of steep bathymetric variation, such as the shelf break. An example is the Bay of Biscay, where the internal tides are the most energetic (Baines, 1982). Internal tides travel slowly compared with surface gravity waves. Hyder et al. (2005) made observations of internal solitons that occurred between January and April 1998 at a water depth of 440 m northeast of the Andaman Islands, Bay of Bengal. Their observations indicated the occurrence of internal solitons with thermocline depressions of up to 50 m and an upper-layer current velocity of up to 120 cm s21. These solitons only occurred during spring tides, when the tidal range exceeded 1.5 m. In the Suruga Trough in Japan, semidiurnal tidal fluctuations are evident in the current with the total amplitude reaching 50 cm s21 at a depth of 1,370 m. These currents have been associated with internal tides (Matsuyama et al., 1993). Velocity measurements associated with internal tides in the Gaoping Submarine Canyon off the Southwest Taiwan have revealed maximum velocities of over 100 cm s21 (Lee et al., 2009). At these velocities, even gravelgrade grains can be eroded and transported by baroclinic tidal currents.

4.8.4 M2 Tidal Energy Dissipation in the Deep Ocean The advances in tidal mapping afforded by the Topex/Poseidon Satellite have allowed Egbert and Ray (2000) to answer some long-standing questions about tidal energetics. In addressing the tidal issues, Egbert and Ray (2000) focused on the principal lunar semidiurnal M2 tide because it accounts for approximately two-thirds of the total planetary dissipation (Cartwright and Ray, 1991). Analyses of the altimeter-derived cotidal charts reveal that although most M2 tidal energy dissipates in shallow seas, about 1 TW (terawatt, 1 TW=1012 W) or 25 30% of the total energy dissipates near rugged bottom topography, such as seamounts and mid-ocean ridges, in the deep

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FIGURE 4.73 Map, based on Topex/Poseidon Satellite altimeter data, showing M2 tidal energy dissipation in the world’s oceans. Note areas of high-energy dissipation (red) in the open ocean (deep water) due to seafloor irregularities. See Egbert and Ray (2000) for details. Image credit: Richard Ray/Space Geodesy branch, NASA/GSFC. NASA Goddard Space Flight Center, http://earthobservatory.nasa.gov/IOTD/view.php?id=654. Accessed August 5, 2011.

open ocean (Figure 4.73). An important implication of this finding is that the tidal energy could be used to explain as possible energy sources needed to maintain the ocean’s large-scale “conveyor-belt” THC and to mix upper ocean heat into the abyssal depths (see Section 4.2). Of the 2 TW required for this ocean-mixing process, about 1 TW is believed to be supplied by the winds. There has been speculation that the tides (Munk and Wunsch, 1998), by pumping energy into vertical water motions, supply the remainder (Egbert and Ray, 2000). The distribution of M2 (semidiurnal) barotropic tidal constituent in the world’s oceans clearly shows that both shallow seas and deep oceans are affected (Figure 4.74, see color plate). In discussing the importance of this altimeter data, Jayne et al. (2004, p. 73) state that “The ocean tides flowing over rough topography create internal waves, some of which break up into turbulence, causing ocean mixing. This mixing determines the stratification of the ocean and influences the ocean’s circulation and heat transport, as well as the penetration of heat and gases from the atmosphere into the deep ocean.” Although how much of the barotropic tidal energy converts into internal waves and tides is still a mystery (Ray et al., 2005), St. Laurent et al. (2003) reported that internal tides are generated commonly at regions where the barotropic tidal current encounters variations in bottom topography. Semidiurnal internal tidal currents are likely major factors in shaping continental slopes. Cacchione et al. (2002) have discussed the intensification of near-bottom water velocities, caused by reflection of semidiurnal internal

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Amphidromic ponits where tidal amplitude is zero

60ºN

30ºN

0º

30ºS

60ºS 60ºE

120ºE 0

180º

120ºE

60ºW

0º

10 20 30 40 50 60 70 80 90 100 110 120 130 cm

FIGURE 4.74 The M2 (semidiurnal) barotropic tidal constituent in the world’s oceans based on Topex/Poseidon satellite altimeter data. Amplitude is indicated by color bar scale. Note high amplitude locations (red-yellow) in certain parts of the world’s continental margins. The white lines are co-tidal differing by 1 hour. The curved arcs around the amphidromic points show the direction of the tides, each indicating a synchronized 6 hour period. See Egbert and Ray (2000) for further details. Image credit: Richard Ray/Space Geodesy branch, NASA/GSFC. NASA  Goddard Space Flight Center. , http://svs.gsfc.nasa.gov/stories/topex/tides.html . Accessed October 8, 2011.

tides, and their role on sedimentation patterns and bottom gradients of continental slopes off northern California and New Jersey. The Indonesian seas serve as the natural laboratory to investigate the role of thermocline (Gordon, 2005) and internal tides (Robertson and Ffield, 2005).

4.8.5 M2 Barotropic and Baroclinic Tides in the Indonesian Seas Oceanographic studies have shown that the Makassar Strait in the Indonesian seas is dominated by strong tidal currents, internal waves/tides, and solitons (Ray et al., 2005). The Indonesian throughflow (ITF), which passes through the Makassar Strait (Gordon, 2005), is a series of ocean currents that flow from the tropical Pacific Ocean through the Indonesian seas into the Indian Ocean (Figure 4.75A, see color plate). It transports nearly 10 Sv (1 Sv or Sverdrup=106 m3 s21) of the Pacific Ocean water into the Indian Ocean. The ITF, a thermocline flow, is stratified along the Makassar sill depth of

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FIGURE 4.75 (A) Map showing ITF pathways and estimates of total volume transport in Sv (Sverdrup 5 106 m3 s21). (B) Panel showing the partitioning of transport above and below the Makassar Strait sill depth of approximately 680 m. After Gordon (2005), with permission from Oceanography.

approximately 680 m (Figure 4.75B, see color plate). Gordon et al. (2010) reported transport values of 11.6 Sv, obtained by the International Nusantara Stratification and Transport (INSTANT) program, for the ITF in the Makassar Strait (Figure 4.76). Current-meter measurements from two moorings in the Labani Channel recorded velocities in excess of 50 cm s21 at 250 m in the Makassar Strait (Wajsowicz et al., 2003). The Indonesian archipelago provides a challenging opportunity to study the dissipation of M2 tidal energy because of the complex bathymetry and the stratified waters. In describing the tidal phenomenon in the Indonesian seas, Ray et al. (2005) state that “Tidal phenomena in the Indonesian seas are among the most complex in the world. Complicated coastal geometries with narrow straits and myriad small islands, rugged bottom topography next to wide shelves of shallow water, and large quantities of tidal power

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112ºE

116ºE

120ºE

124ºE

128ºE

132ºE

4ºN

8ºN

Borneo

0º

0º

4ºs

4ºS

8ºs

8ºS

12ºs

12ºS 108ºE

112ºE

116ºE

120ºE

124ºE

128ºE

132ºE

FIGURE 4.76 Map showing complex bathymetry of the Indonesian seas. The map also shows transport values in 106 m3 s21 within the passages measured by the INSTANT program, 2004 2006. The italics numbers in black represent transport values based on pre-INSTANT data. The red numbers are the 2004 2006 3-year mean transports measured by INSTANT. After Gordon et al. (2010), with permission from Elsevier.

input from the adjoining Indian and Pacific Oceans—all combine to form a complex system of interfering three-dimensional waves.” Ray et al. (2005) used the Topex/Poseidon Satellite altimetry mapping of the Indonesian seas (Figure 4.77) in evaluating the role of M2 barotropic tides. Altimeter-derived cotidal chart (Figure 4.78A) clearly reveals that the semidiurnal response is dominated by the large tide from the Indian Ocean, where amplitudes are well over a meter off northwest Australia. This wave passes into the Banda and Flores Seas. From the Banda Sea, the semidiurnal tide passes slowly northward through the Molucca Sea region. From the Flores Sea, it propagates slowly northward into Makassar Strait and also westward, but more weakly, across the Java Sea (Ray et al., 2005). The M2 barotropic tidal current velocities commonly reflect local bathymetry, with shallow water giving rise to high-velocity currents (Figure 4.78B). In contrast to the study of M2 barotropic tides by Ray et al. (2005), Robertson and Ffield (2005) have focused on the generation and propagation of M2 baroclinic tides in the Indonesian seas. Higher amplitudes occur in the Sulawesi Sea, Indian Ocean, and shallow regions, including along the Australian coast where amplitudes exceed 150 cm (Figure 4.79A). Lower amplitudes, ,20 cm, occur in the Java and Ceram Seas, and in the Timor Sea. The M2 tide enters the Indonesian seas both from the Indian Ocean and the Pacific Ocean, with stronger tides from the Indian Ocean. As revealed by

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FIGURE 4.77 Map of the Indonesian seas showing the ground tracks (red lines) of the Topex/ Poseidon Satellite. After Ray et al. (2005), with permission from Oceanography.

the phase lag diagram (Figure 4.79B), the M2 tide progresses from the Indian Ocean through the Timor Sea, where an amphidromic point forms. And, a portion passes through Lombok Strait into the region south of Makassar Strait. In the Timor Sea, it splits with one portion progressing into the Banda Sea and another into the Arafura Sea. Robertson and Ffield (2005) have demonstrated the importance of rough bottom topography in generating M2 internal tides. Transects of the major axes of the M2 tidal ellipses from the depth-dependent velocity for the Lifamatola Strait, Timor and Ombai Straits, Maluku Strait, and Makassar Strait reveal baroclinic responses (Figure 4.80). Along the transect in the Timor Sea and across Ombai Strait into the Banda Sea (see Figure 4.76 for locations), baroclinic tides are generated primarily in three locations: (1) over the Australian slope, (2) on the southern shoulder of Timor Island propagating offshore, and (3) over the sill north of the Ombai Strait propagating both into the strait and northwest into the Banda Sea (Figure 4.80). The strongest internal tides occur in the Ombai Strait, with major axes .50 cm s21. Figure 4.80C shows that in the Maluku Sea, internal tides generated over rough topography propagate into the upper water column, where the internal wave rays become more horizontal due to stronger stratification. The transect through the Makassar Strait shows a strong baroclinic tide

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FIGURE 4.78 (A) Cotidal chart of the semidiurnal tide M2 in the Indonesian seas. Amplitude of the M2 barotropic tide is shown by the color bar scale. (B) Maximum barotropic current velocity of M2 tide. Note nonlinear color bar. Velocity units are cm s21. In general, these velocities reflect bathymetry, with shallow water giving rise to rapid currents. Currents in the Banda, Timor, and part of the Flores Seas tend to rotate anticlockwise; currents elsewhere are generally close to rectilinear, although the deep Pacific currents rotate clockwise. After Ray et al. (2005), with permission from Oceanography.

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FIGURE 4.79 (A) The elevation amplitude for the M2 baroclinic tidal constituent over the model domain. (B) The phase lag for the M2 elevation over the model domain with the direction of propagation indicated by white arrows. The tide propagates from the Indian Ocean through the Timor Sea, where an amphidromic point forms. A portion of the tide also propagates from the Indian Ocean through Lombok Strait into the region south of Makassar Strait. In the Timor Sea, the tide splits with a portion going into the Arafura Sea and another into the Banda Sea. In the Banda Sea, the tide splits again with a portion going into the Flores Sea and another north into the Ceram Sea. South of Makassar Strait is a confluence region for the portions from Lombok Strait, the Flores Sea, and the weaker portion from the Pacific Ocean (thinner arrows) via Makassar Strait. From there, it exits via the Java Sea. After Robertson and Ffield (2005), with permission from Oceanography.

originating at the western shoulder and propagating across the strait and a bottom-trapped tide at the eastern edge (Figure 4.80D). The Dewankang Sill at the southern end of Makassar Strait is also known for generating baroclinic tides due to topographic irregularities (Hatayama, 2004). In conclusion, strong, baroclinic M2 tides were generated along the shelf slope break and over rough topography, particularly in straits in the Indonesian seas. Importantly, the baroclinic tidal velocities exceeded the barotropic velocities over most of the domain. New data obtained by the INSTANT program validate the above conclusions (Robertson, 2010). Despite these advances in oceanography of baroclinic currents, our understanding of sedimentological aspects of baroclinic currents is zilch.

4.8.6 Sedimentological Criteria The potential significance of shoaling internal waves for causing sediment movement on continental shelves and slopes has been discussed by Cacchione and Southard (1974). Laboratory experiments confirmed that shoaling interfacial waves could generate ripples and larger bedforms in artificial sediment (Southard and Cacchione, 1972). Stride and Tucker (1960) attributed the development of modern sand waves near the shelf edge to internal waves. Karl et al. (1986), using sparker profiles, documented sand

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(A)

(B)

(C)

(D)

Distance (km) FIGURE 4.80 Transects of the major axes of the M2 (baroclinic) tidal ellipses from the depthdependent velocity for the (A) Lifamatola Strait, (B) Timor and Ombai Straits, (C) Maluku Strait, and (D) Makassar Strait. Large baroclinic responses occur in all transects. Velocity values (side bar) are in cm s21. See Figure 4.75 for locations. Compiled from Robertson and Ffield (2005), with permission from Oceanography.

waves in the heads of submarine canyons of the Bering Sea. In this case, a surface sediment sample (C1) was composed of 19% gravel, 76% sand, and 5% mud. The modal class of this sample was fine sand. However, no sedimentary structures were described from the cores of these sand waves. Karl et al. (1986) speculated that internal waves were responsible for the origin of sand waves. They also suggested that delivery of large volumes of freshwater and large quantities of sediment directly to the heads of submarine canyons during periods of low sea level might have enhanced the propagation of high-frequency internal waves. Gao et al. (1998) interpreted ancient strata with bidirectional cross-bedding (see also Gao and Eriksson, 1991), flaser bedding, wavy bedding, and

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lenticular bedding as deposits of internal tides based on associated deep-water turbidite and slump facies. The supreme evidence for interpreting deposits of “internal tides” or baroclinic currents in the rock record is the proof for tidal currents in a stratified deep ocean. Without that evidence for density stratification (pycnocline), there is no difference between a tidal deposit formed by surface (barotropic) tide in a shallow-marine shelf and a tidal deposit formed by internal (baroclinic) tide in a deep-marine slope or canyon environment. Furthermore, not all tidal bottom currents in submarine canyons are related to density stratification. Unlike barotropic tidal currents that flow along the axis of the canyon, baroclinic currents flow across the canyon and in the direction parallel to the shelf break (Allen and Durrieu de Madron, 2009). If so, these two different types of currents may not generate the same kind of deposits. For example, the origin of bidirectional cross-bedding by internal tides in submarine channels (Gao and Eriksson, 1991) would be difficult to explain if baroclinic currents were to flow across the channel, instead of up- and down- the channel. He et al. (2011, their Figure 11) recently proposed the first facies model (i.e., ideal vertical sequence) for internal tide-related deposits. This sequence is composed of a basal sandy turbidite division (Layer I), a middle sandy division with traction structures formed by reworking by internal waves and internal tides (Layer II), and an upper hemipelagic muddy division (Layer III). This facies model is untenable for the following reasons (Shanmugam 2011e). 1. The ideal sequence, in a bizarre way, mimics Ta, Tc, and Te divisions of the “Bouma Sequence”. But the authors do not make any reference to the original turbidite facies model of Bouma (1962). 2. The basal massive sandstone layer has been interpreted as turbidites. But the origin of massive sands can be explained by numerous alternative deepwater processes: (a) contour currents (Rodriguez and Anderson, 2004), (b) bed-load deposition (Sanders, 1965), (c) grain flows (Stauffer, 1967), (d) pseudoplastic quick bed (Middleton, 1967), (e) density-modified grain flows (Lowe, 1976), and (f) sandy debris flows (Shanmugam, 1996a). 3. He et al. (2011) reasoned that bidirectional cross-bedding cannot be explained by either turbidity currents or by contour currents, and therefore it must be formed by internal tides. But bidirectional cross-laminae have been documented in deep-water traction deposits associated with wind-driven Loop Current in the Gulf of Mexico (Shanmugam et al., 1993a, 1993b, 1995c). The Loop Current generates eddies that can create complex current orientations. 4. He et al. (2011, their Figure 2) proposed that their study area is located in an open “abyssal basin” environment. But internal waves and tides are commonly associated with shelf-edge settings (Inman et al., 1976, see their Figure 4). Although it is conceivable that weak internal tides may be generated near seamounts in the open ocean, seamounts are ineffective barriers for generating major internal tides (Holloway and Merrifield,

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1999). Importantly, there is no evidence of seamount near the study area in Figure 2 of He et al. (2011). Finally, internal tides are akin to tsunamis in terms of our ignorance. Analogous to tsunamis (Shanmugam, 2011c), we know nothing about the flow behavior of internal tides. The implication here is that the less the knowledge we possess on a given oceanographic phenomenon, the more the nonsense we propagate on its purported deposits. In order to interpret a sedimentary feature as the product of internal tides, one needs to establish its association with paleo pycnoclines. Thus far, sedimentologists who study the phenomenon of internal tides have realized that it is impossible to provide physical evidence for (1) density-stratified waters and (2) for the timing and depth at which the discussed sedimentary features were developed. Until we develop objective and reliable sedimentological criteria for recognizing deposits of baroclinic currents (internal tidalites) of stratified water bodies, it is preferred to classify deep-water deposits with tidal signatures as products of “deep-water tidal bottom currents” rather than that of “baroclinic currents” (i.e., internal tidalites). This is a topic for future research on deep-water tidal sedimentation.

4.9 PROBLEMATIC BEDFORM-VELOCITY MATRIX FOR DEEPWATER BOTTOM CURRENTS Stow et al. (2009) proposed a bedform-velocity matrix for deep-water bottom currents (Figure 4.81). This matrix diagram is a slightly modified version of Figures 3.1 and 3.2 in Belderson et al. (1982). There are fundamental problems with this matrix diagram. 1. The concept of bedform-velocity matrix became popular in the 1960s with the advent of matrix diagrams of alluvial sedimentary structures based on empirical data derived from flume experiments (Simons et al., 1965; see also Southard, 1975; Southard and Boguchwal, 1990). However, the matrix diagram proposed by Stow et al. (2009) is not based on empirical data from experiments. 2. There are at least four different types of deep-water bottom currents, namely (1) thermohaline-induced geostrophic bottom currents (i.e., contour currents), (2) wind-driven bottom currents, (3) deep-water tidal bottom currents, and (4) baroclinic currents (internal tides). Because each bottom-current type behaves differently, it is inappropriate to develop a single matrix diagram as if all four types of bottom currents behave the same. 3. Belderson et al. (1982) established their bedform-velocity matrix based on years of research on natural bedforms developed by tidal currents on the shallow-water shelf (,200 m water depth) areas of Bristol Channel, English Channel, Celtic Sea, etc. They systematically documented their results in several chapters in a book (edited by Stride, 1982) entitled “Offshore Tidal Sands: Processes and Deposits.” Chapter 3 in that book,

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FIGURE 4.81 Bedform-velocity matrix for deep-water bottom currents. After Stow et al. (2009), Geological Society of America.

written by Belderson et al. (1982), is the original source of data for the matrix diagram of Stow et al. (2009). 4. Stow et al. (2009) applied the bedform-velocity matrix, developed by Belderson et al. (1982) for shelf tidal currents, to all four types of deep-water bottom currents. This logic is fundamentally flawed for the following reasons: (a) Shallow-water tidal currents and four types of deep-water bottom currents are not one and the same hydrodynamically. (b) Shallow-water tidal currents are bidirectional, whereas deep-water contour currents, for example, are unidirectional. (c) Stow et al. (2009) acknowledged that although the velocity data presented by them are for near-bottom flow, they did not define the exact height above seafloor.

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(d) Stow et al. (2009) conceded that they did not address the variable nature of the benthic boundary layer that will also complicate how flow velocity affects seafloor bedform. (e) Stow et al. (2009) acknowledged that for most of their data sets it is impossible to know the precise flow velocity (mean or peak) that created the observed bedform, which is the most important factor of all. (f) Stow et al. (2009) conceded that they rarely had the opportunity of witnessing the development of deep-water bedforms in situ. (g) Stow et al. (2009) also acknowledged that they did not consider the effects of sediment supply and bed roughness on bedform development. (h) Unlike Belderson et al. (1982) who gathered all their bedform and velocity data themselves, Stow et al. (2009) mostly relied on published

Golden Gate Bridge

Depth: 113 m

Tidal sand waves

FIGURE 4.82 Image showing a field of giant tidal sand waves and other bedforms at the mouth of San Francisco Bay. The view is from the northwest toward the Golden Gate Bridge (seen in the background), which is approximately 2-km long. More than 40 large (greater than 50-m wavelength) sand waves were mapped, with crest-to-crest lengths of as much as 220 m (722 ft) and heights of as much as 10 m (33 ft). This computer-generated image by Patrick Barnard of sand waves is based on high-resolution multibeam imaging of the seafloor using research vessel VenTresca by the CSUMB Seafloor Mapping Laboratory. Vertical exaggeration: 3X. Geological details of the setting and sand waves are discussed by Barnard et al. (2006). The land was imaged using digital orthophotos draped over a U.S. Geological Survey digital elevation model. The Golden Gate Bridge model is courtesy of IVS 3D©. Image courtesy of USGS, http://soundwaves.usgs.gov/2006/09/ViewtoGateHGLG.jpg. Accessed July 14, 2011.

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data gathered by others. Although this is an acceptable practice, it does allow for inconsistencies between authors. (i) Finally, Belderson et al. (1982) documented tidal sand waves on the continental shelf. In fact, some of the world’s largest tidal sand waves have been mapped by Barnard et al. (2006) at the mouth of San Francisco Bay (Figure 4.82, see color plate). These sand waves, composed of gravel and coarse sand, have an average wavelength of 82 m and an average height of 6 m. They are formed by ebb tidal currents with a velocity of 2.5 m s21 in water depths ranging between 35 m and 80 m. To date, no such giant sand waves deposited by tidal currents have been documented in deep-water environments. In the absence of such real-world examples, it is inappropriate to advocate a unified bedform-velocity matrix for all four types of deep-water bottom currents. On one hand, Stow and Fauge`res (2008) advocate that interpretation of bottom currents can be made on the basis of bioturbation and grain-size variations alone using their contourite facies model (Figure 4.19), without the benefit of primary structures, and on the other hand, Stow et al. (2009) advocate that the primary sedimentary structures are the key for interpreting bottom currents based on their bedform-velocity matrix (Figure 4.81). The apparent incongruity between the model (Figure 4.19) and the matrix (Figure 4.81) is truly remarkable!

4.10 PROBLEMS WITH INTERPRETATION OF SEISMIC FACIES AND GEOMETRIES 4.10.1 Seismic Facies In their comprehensive review of seismic expression of contourite depositional systems, Nielsen et al. (2008) state that “... because the reflections result from changes in the physical parameters through the sedimentary succession, there is no unequivocal correlation between seismic facies and sedimentary structures within the facies. A seismic facies characterized by a parallel reflection configuration, for example, need not necessarily indicate the existence of fine parallel banding or stratification of the sediments.” Clearly, there are fundamental problems in using seismic facies for interpreting bottom-current deposits, given the fact that BCRS are difficult to recognize even from the direct examination of the rocks. Major practical problems are as follows: 1. Although common seismic reflections of contourites are (a) continuous and parallel, (b) wavy, and (c) transparent (Nielsen et al., 2008), these seismic facies are also found in MTDs (Figure 3.79). 2. Parallel reflection pattern is not unique to contourites. Parallel reflections could also be interpreted as either basin-floor turbidites or hemipelagites.

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3. Parallel reflections reveal nothing unique about the type of bottom currents, such as contour-following currents, wind-driven eddies, or up- and downflowing tidal currents in submarine canyons.

4.10.2 Wave Geometry Bottom-current generated sediment waves have been documented in the world’s oceans that include the following: 1. Northern Rockall Trough, Northwest Atlantic (Howe et al., 1994; Masson et al., 2002) 2. Blake-Bahama Outer Ridge, North Atlantic (Flood, 1994) 3. Gulf of Mexico (Kenyon et al., 2002b) 4. Mediterranean Sea (Kenyon and Belderson, 1973) 5. Southwestern Adriatic Margin (Verdicchio et al., 2007) 6. Argentine Basin (Klaus and Ledbetter, 1988) 7. Antarctica (Cunningham and Barker, 1996; Rebesco et al., 2007), among others. In the northern Rockall Trough (Figure 4.83), bottom-current activities were studied by Masson et al. (2002). They recognized two water masses, namely an upper Eastern North Atlantic Water (ENAW) and a lower NADW. Current velocities of 10 50 cm s21 are common for ENAW. On the southern flank of the Wyville Thomson Ridge (Figure 4.83), at a 477 m water depth, currents of up to 100 cm s21 were measured. Masson et al. (2002) documented four types of sediment features related to bottom-current activities: (1) broad-sheeted drifts (Figure 4.84), (2) elongate drifts (Figure 4.85A), (3) sediment waves (Figure 4.85B), and (4) thin contourite sheets. Sediment waves with wavelength of 1 2 km and wave height of 20 m are present. Sediment waves are composed mostly of mud (Howe, 1995). The exact processes that initiate sediment waves are still poorly understood (Wynn and Masson, 2008, p. 292). Well-developed wave geometries, interpreted as mega sediment waves formed by the Mediterranean outflow water off Southwest Portugal (Figure 4.86), have been reported (Nielsen et al., 2008). Seismic wave geometries induced by large-scale seafloor features (10 80 m in height), such as “migrating mud waves” or “abyssal bedforms,” have been reported from the modern oceans (Flood, 1988; Klaus and Ledbetter, 1988). Mud waves were ascribed to sculpting of muddy seafloor by deep bottom currents, such as the AABW in the Argentine Basin (Klaus and Ledbetter, 1988). Nevertheless, seismic wave geometry of modern contourite deposits is of no relevance for interpreting ancient muddy contourite deposits in the stratigraphic record. This is because: 1. In the Southwestern Adriatic Margin, sediment waves have been documented on high-resolution seismic profiles. These sediment waves are

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FIGURE 4.83 Location map showing northern Rockall Trough area used in studying bottomcurrent sediment waves, Northwest of Scotland. Filled arrows, surface circulation; open arrows, deep-water circulation. After Masson et al. (2002), with permission from Elsevier.

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FIGURE 4.84 A 3.5 kHz seismic profile showing broad-sheeted drift. Note large-scale asymmetry of sediment units across the drift crest and sediment waves on the southwest flank of the drift. Northern Rockall Trough. After Masson et al. (2002), with permission from Elsevier.

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FIGURE 4.85 (A) Deep-tow boomer profile showing upslope-migrating sediment waves (boxed area) and elongate drift. (B) High-resolution seismic profile showing enlarged version of upslope-migrating sediment waves (see boxed area in A). More sites of erosion and deposition. Compiled from Masson et al. (2002), with permission from Elsevier. See also Howe et al. (1994) for additional details.

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FIGURE 4.86 Single-channel high-resolution seismic profile showing wave geometry, interpreted as mega sediment waves, off Southwest Portugal. After Nielsen et al. (2008), with permission from Elsevier.

characterized by “...layered, continuous, convex upwards, subparallel reflectors...” (Verdicchio et al., 2007, p. 205 and their Figure 3). The importance of this study is that it has cored sediment intervals from the upper 15 m of the sediment wave (Verdicchio et al., 2007, their

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Figure 10). These authors (p. 209) describe the cored sediment as follows: “The most common lithology recovered from the drift is Foraminifera-rich intensely bioturbated mud or silt-rich mud. . . . Primary sedimentary structures related to traction processes have not been observed.” The equivalent of this cored sediment facies in the ancient record would simply be a foraminiferal, bioturbated mudstone, or shale without any primary sedimentary structures. Such a mudstone/shale interval cannot be interpreted as deposits of bottom currents or contour currents, without primary sedimentary structures and other information on regional setting and paleo-ocean circulation patterns. Physical diagenesis of biogenic siliceous deep-sea sediments is an important factor in understanding compaction and porosity reduction (Hesse and Schacht, 2011) that has implications for seismic geometry. For example, a sharp porosity reduction from 75% to 35% occurs over a few tens of meters of burial during conversion of opal-A to opal-CT (Tada, 1991). In general, a marine mud unit (with sediment wave geometry) would undergo nearly 50% reduction in volume between deposition and deep burial due to compaction. As a consequence, the density of mud would nearly double with burial. This increase in density would affect the Pwave velocity (Gardner et al., 1974; Sheriff and Geldart, 1995). However, the influence of diagenesis (i.e., compaction and cementation) on seismic reflection patterns of the “muddy wave geometry” in the rock record remains an unresolved issue. The other issue is distinguishing between “mud wave” and “sand wave” in the rock record using seismic geometry. Kenyon et al. (2002b) reported sand wave fields beneath the Loop Current in the Gulf of Mexico. The term “sand wave” has a specific meaning in terms of grain size (.0.2 mm, fine sand) and flow velocity (.40 cm s21) in sedimentology (Southard, 1975, his figures 2-5). Therefore, the casual usage of the term “sediment wave,” for seismic features of contourites, without knowing their grain size and flow velocity, could be misleading. This is because (a) the term “sediment wave” could mean either “mud wave” or “sand wave,” and (b) contourites have both muddy and sandy types (Stow and Lovell, 1979). Deepwater (muddy) wave geometries should not be confused with dune bedforms in rivers that create cross-bedding due to bed-load transport of granular material, which must be larger than 125 µm in grain size (i.e., fine sand). Deep-sea migrating waves are composed primarily of mud (i.e., silt and clay), and therefore they do not have the necessary sand grade (i.e., granular material) to generate cross-bedding. Giant sediment waves (5 m in height) on the continental margin off Nice (southern France) that are composed of sand and boulders were ascribed to deposition by “sediment flows” (Malinverno et al., 1988). These sediment flows have been considered to be a combination of both debris flow

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and turbidity current (Malinverno et al., 1988). But it is unclear as to how these sediment waves were deposited by two rheologically different sediment flows composed of plastic debris flows and Newtonian turbidity currents. 6. Seismic wave geometries have been associated with (a) deep-water muddy contourites (Nielsen et al., 2008), (b) shallow-water muddy contourites (Verdicchio and Trincardi, 2008), (c) shallow-water sandy tidalites (Belderson et al., 1982), and (d) deep-water turbidites (Normark et al., 1980). However, there are no objective criteria to distinguish wave geometry created by contour currents from wave geometry created by tidal currents or by turbidity currents on seismic profiles. 7. Wynn and Stow (2002) proposed four criteria to distinguish sediment waves deposited by bottom currents: (a) depositional site away from turbidity-current input, (b) absence of coarsening up- or downslope trend in wave dimensions, (c) contourites with intense bioturbation, (d) wave crest aligned oblique to bathymetric contours, and (e) absence of consistent spatial trend in sediment-wave sequence thickness. Given the controversies that surround some of these criteria (e.g., bitoturbation, see Section 4.5.4), these criteria are difficult to apply even to modern examples. These criteria are most certainly impractical to apply to the ancient record.

4.10.3 Channel-Levee Geometry Rebesco et al. (2007, their Figure 1) discussed the similarity in seismic geometry between turbidite channel-levee systems and contourite drifts. The difference is that contourite drifts show asymmetric moat and mound geometry, whereas turbidites exhibit symmetric gull-wing geometry. But in the rock record, it would be difficult to distinguish one from the other on seismic profiles. Seismic channel geometry (Figure 4.87, see color plate) has been reported for the Santos Channel, a 100 km-long channel-like gutter at the foot of an intraslope escarpment, in the Santos basin in offshore Brazil by Duarte and Viana (2007). This channel was excavated by the strong northwardflowing Southern Ocean Current (SOC) during early Miocene. Analogous to the present-day AABW, the SOC was an along-slope flowing contour current. The SOC not only formed the channel by erosion but also deposited the adjacent contourite drifts on both sides. The Santos Channel and adjacent contourite drifts (Figure 4.87, see color plate) resemble those of turbidite channels with adjacent levee complexes. In the stratigraphic record, distinguishing contourite channels that align parallel to the slope from turbidite channels that generally align perpendicular to the regional slope is a challenge. As mentioned earlier, turbidity currents may flow parallel to the strike of the regional slope along trench floors (Underwood and Bachman, 1982).

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FIGURE 4.87 Seismic block diagram showing channel geometry of the Santos Channel caused by the erosion of seafloor by the along-slope flowing (northward) SOC during early Miocene. The paleocurrent direction is same as the present-day direction of the AABW. Cross-sectional channel geometry in seismic profiles could be misinterpreted as turbidite channels in other areas. Note levee-like geometry on both sides of the channel. Also note continuous and parallel seismic reflections of contourite deposits on the right-hand side showing sheet geometry. After Duarte and Viana (2007), with permission from the Geological Society of London.

FIGURE 4.88 Seismic profile showing sheet geometry with continuous and parallel reflection patterns. Note position of cored wells through two sand units examined in the Ewing Bank area, Gulf of Mexico. Core from the L-1 sand shows a dominance of BCRS (e.g., Figures 4.26 and 4.27). After Shanmugam et al. (1993a), reprinted by permission of the American Association of Petroleum Geologists whose permission is required for further use.

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4.10.4 Sheet Geometry In the Gulf of Mexico, a seismic profile shows sheet geometry (Figure 4.88) continuous and parallel reflection patterns for the L-1 Sand (Upper Pliocene) that has been interpreted to be composed dominantly of deposits of the wind-driven Loop Current in the Ewing Bank area (Shanmugam et al., 1993a). Our interpretation was based on integration of core details with regional current pattern in the Gulf of Mexico. In the absence of core calibration, one could misinterpret the L-1 Sand, showing sheet geometries (Figure 4.88), as sheet-like turbidites or hemipelagites. The oceanographic details on the type of bottom current cannot be inferred from the nature of seismic reflection patterns (e.g., parallel reflections). Sheet geometries with continuous and parallel reflections, for example, are associated with contourite deposits in offshore Brazil (Figure 4.87, see color plate) and in offshore Norway (see Figure 3.80, see color plate).

4.11 SYNOPSIS The three vertical segments of the world’s oceans, composed of (1) the upper surface currents, (2) the middle deep-water masses, and (3) the lower bottom currents, form a vertical continuum. Deep-water bottom currents are products of thermohaline, wind, and tidal forces. A distinctive attribute of BCRS is their traction structures. Double mud layers and sigmoidal cross-bedding are unique to deep-water tidal deposits in submarine canyons. The contourite facies model and the bedform-velocity matrix for deep-water bottom currents are unsustainable. Seismic facies and geometries are unreliable for distinguishing individual types of BCRS in the ancient record.