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Modern internal waves and
internal tides along oceanic
pycnoclines: Challenges and
implications for ancient
deep-marine baroclinic sands
G. Shanmugam
ABSTRACT
Thus far, the subject of deep-marine sands emplaced by baro-
clinic currents associated with internal waves and internal tides
as potential reservoirs has remained an alien topic in petro-
leum exploration. Internal waves are gravity waves that oscil-
late along oceanic pycnoclines. Internal tides are internal waves
with a tidal frequency. Internal solitary waves (i.e., solitons),
the most common type, are commonly generated near the shelf
edge (100–200 m [328–656 ft] in bathymetry) and in the deep
ocean over areas of sea-floor irregularities, such as mid-ocean
ridges,seamounts,andguyots.Empirical data from 51 locations
in the Atlantic, Pacific, Indian, Arctic, and Antarctic oceans re-
veal that internal solitary waves travel in packets. Internal waves
commonly exhibit (1) higher wave amplitudes (5–50 m [16–
164 ft]) than surface waves (<2 m [6.56 ft]), (2) longer wave-
lengths (0.5–15 km [0.31–9mi])thansurfacewaves(100m
[328 ft]), (3) longer wave periods (5–50 min) than surface
waves (9–10 s), and (4) higher wave speeds (0.5–2ms
–1
[1.64–6.56 ft s
–1
]) than surface waves (25 cm s
–1
[10 in. s
–1
]).
Maximum speeds of 48 cm s
–1
(19 in. s
–1
) for baroclinic cur-
rents were measured on guyots. However, core-based sedi-
mentologic studies of modern sediments emplaced by baroclinic
currents on continental slopes, in submarine canyons, and on
submarine guyots are lacking. No cogent sedimentologic or
seismic criteria exist for distinguishing ancient counterparts.
Outcrop-based facies models of these deposits are untenable.
Therefore, potential exists for misinterpreting deep-marine
baroclinic sands as turbidites, contourites, basin-floor fans, and
AUTHOR
G. Shanmugam Department of Earth and En-
vironmental Sciences, University of Texas at Arlington,
Arlington, Texas 76019; shanshanmugam@aol.com
G. Shanmugam is a pragmatic deep-water process
sedimentologist. He obtained his M.Sc. degree in
applied geology (1968) from the Indian Institute of
Technology in Bombay and his Ph.D. in geology (1978)
from the University of Tennessee at Knoxville. He
worked for Mobil in Dallas, Texas, from 1978 to 2000.
HeismarriedtoJean(1976–present). He organized
workshops on deep-water sandstone petroleum
reservoirs, using conventional cores, in the United
States, United Kingdom, Brazil, China, and India. His
current research includes sandy mass-transport pro-
cesses, bottom currents, and oceanic waves (internal,
cyclonic, and tsunami). He has 330 published works,
including two volumes of Elsevier’sHandbook of Pet-
roleum Exploration and Production (2006 and 2012).
ACKNOWLEDGEMENTS
I thank C. R. Jackson of Global Ocean Associates for
providing an updated figure on the locations of
observed internal waves (Figure 1). I also thank
Elsevier, Springer, Academic Press (Elsevier),
Copyright Clearance Center’s RightsLink, National
Aeronautics and Space Administration, the Ocean-
ography Society, AAPG, and the Geological Society of
America for granting permission to reproduce figures. I
thank AAPG Editor Stephen E. Laubach (Bureau of
Economic Geology, Austin, Texas) and associate editor
William A. Hill (British Petroleum, Houston, Texas) for
their helpful comments. The selection of David J. W.
Piper (Geological Survey of Canada and Bedford
Institute of Oceanography, Halifax, Nova Scotia), one
of the three editors-in-chief of Marine Geology,as
the principal reviewer by AAPG and his willingness to
review this manuscript, which deals with physical
oceanography and process sedimentology, is much
appreciated. I also thank AAPG technical editor An-
drea Sharrer, consulting editor Carol Christopher,
technical publications coordinator Paula Sillman, and
AAPG Bulletin consultant Frances Plants Whitehurst
for their help during the various stages of progress of
this paper. As always, I thank my wife, Jean, for her
general comments. This paper is the culmination of
gathering and analyzing empirical data during Jan-
uary 2005 to September 2012 on internal waves,
internal tides, cyclonic waves, and tsunami waves.
The AAPG Editor thanks the following reviewers for
their work on this paper: William A. Hill and David
J. Piper.
Copyright ©2013. The American Association of Petroleum Geologists. All rights reserved.
Manuscript received June 24, 2012; provisional acceptance September 24, 2012; revised manuscript
received October 7, 2012; final acceptance October 17, 2012.
DOI:10.1306/10171212101
AAPG Bulletin, v. 97, no. 5 (May 2013), pp. 799 –843 799
others. Economic risks associated with such mis-
interpretations could be real.
INTRODUCTION
Benjamin Franklin, in 1762, demonstrated that in-
ternal gravity waves on the interface between oil
and water have a much longer period than do sur-
face waves with the same wavelength (Phillips,
1974). Early observations of internal waves in na-
ture have been attributed to Russell (1838), Wallace
(1869), Nansen (1900), and even to earlier Viking
times (Ekman, 1904). Theoretical, numerical, and
observational analyses have presented an exhaustive
and complex account of modern internal waves and
internal tides in the oceans of the world (LaFond,
1962; Wunsch, 1969, 1975; Inman et al., 1976;
Garrett and Munk, 1979; Miles, 1980; Holloway,
1987; Dushaw, 2000; Egbert and Ray, 2000; Jackson
and Apel, 2002; Wunsch and Ferrari, 2004; Simmons
et al., 2004; Apel et al., 2006; Helfrich and Melville,
2006; Garrett and Kunze, 2007; Lavelle and Mohn,
2010; Martini et al., 2011; Alford et al., 2012; Carter
et al., 2012; Hutter, 2012; Jackson et al., 2012; St.
Laurent et al., 2012). However, absolutely no core-
based sedimentologic studies exist on the origin of
primary sedimentary structures in sands emplaced
by baroclinic currents associated with internal
waves and tides in modern marine environments.
An academic divide exists between oceanographic
studies and sedimentologic studies because ocea-
nographers tend to deal with internal waves and
tides that propagate along pycnoclines and that
shoal on continental shelves and slopes, whereas
sedimentologists focus on the emplacement of
sediment by baroclinic currents on the sea floor.
The disparate literature coverage also exists. For
example, Oceanography has just published a spe-
cial issue on internal waves (St. Laurent et al.,
2012), whereas the topic is totally absent in a re-
cent book on Deep-Sea Sediments (Hüeneke and
Mulder, 2011). Despite these empirical and in-
tellectual impediments, a select group of re-
searchers have interpreted ancient strata as depos-
its of internal waves and tides using outcrops (Gao
and Eriksson, 1991; Gao et al., 1998; He et al., 2008,
2011, 2012; Pomar et al., 2012; Bádenas et al.,
2012). These outcrop-based interpretations have
associated limitations (Shanmugam, 2012a, 2013).
From the industry side, an early motivation for
investigating internal waves was that they unex-
pectedly imposed large stresses on offshore oil-
drilling rigs (Apel, 2002). Osborne and Burch (1978)
discussed the importance of studying internal waves
that were observed during drilling by the drillship
Discoverer 534 in the Andaman Sea, offshore Thai-
land, in water depths ranging from 579 m (1899 ft)
to more than 1037 m (3401 ft), for designing pro-
duction facilities in deep water. From a reservoir
viewpoint, no documented examples of petroleum-
producing reservoir sands formed by internal waves
and internal tides exist. A possible exception may be
the Tertiary sands in the Kutei Basin in Indonesia.
These sands were originally interpreted as deep-
water turbidites by Saller et al. (2006), but these
reservoir sands could alternatively be interpreted
as tidalites formed by deep-marine tidal currents
(Shanmugam, 2008a). Such an alternative interpre-
tation merits attention because of the compelling
evidence that exists in the form of (1) documented
oceanic thermoclines (Gordon, 2005), (2) observed
internal waves (Hatayama, 2004; Pujiana et al.,
2009), (3) observed internal tides (Ray et al., 2005),
and (4) measured velocities of deep tidal currents
(Nummedal and Teas, 2001; Wajsowicz et al., 2003)
in the Makassar Strait where the Kutei Basin is
located.
The Kutei Basin example suggests a pressing
need for basic information on internal waves and
internal tides in evaluating deep-marine sands in
petroleum exploration. Therefore, the primary pur-
pose of this review is to amass published oceano-
graphic and sedimentologic facts on internal waves
and tides in building a knowledge base for petro-
leum geologists. In covering the enormity of data
and the diversity of topics, 332 publications (both
print and online) from 1838 to January 2013 have
been cited. Specific objectives are (1) to compile
empirical data on the physical characteristics of
modern internal waves and tides from 51 regions
of the oceans of the world (Figure 1), (2) to pro-
vide a conceptual clarity on sites of deposition of
baroclinic sands in deep-marine environments, (3)
800 Internal Waves, Internal Tides, and Baroclinic Sands
to identify sedimentologic challenges in distinguish-
ing baroclinic sands, and (4) to offer conditional
comments on reservoir geometry and quality.
BASIC CONCEPTS AND NOMENCLATURE
The concept of “baroclinic”is of paramount im-
portance in understanding currents associated with
internal waves and internal tides. According to the
Cooperative Institute for Marine and Atmospheric
Studies (CIMAS, 2012), “Barotropic is the depth-
independent part of the flow. In classic wind-driven
ocean circulation theory, it is the flow that results
from, or is in balance with, a sea surface slope. The
barotropic component of a flow has also been de-
fined as the depth-averaged flow and the flow at the
zero-crossing depth of the first baroclinic mode.
Barotropic instability is the process in which meso-
scale turbulence uses the kinetic energy of the
mean flow to grow. This process occurs in regions
of strong ocean currents like western boundaries
and along the equator. Baroclinic is the depth-
dependent part of the flow. It is the component
of the flow that results from the density distribu-
tion of the fluid. It is the component of flow that
acts to cancel the sea surface flow…. Baroclinic
instability is the process in which mesoscale tur-
bulence uses the available potential energy con-
tained in stratified fluids. This process occurs in
regions with large vertical gradients in buoyancy
due to temperature and salinity differences.”
In general, barotropic flows occur when levels
of constant pressure in the ocean are always par-
allel to the surfaces of constant density, whereas
Figure 1. (A) Map showing the locations (red dots) of observed oceanographic internal waves and tides in coastal seas and in the open
ocean (from Apel, 2002; Jackson, 2004a). (B) Explanation of symbols and numbers. Yellow triangles and numbers represent the locations
of internal waves used for physical properties in this study (Table 2). Depths of pycnoclines in most locations are given in Table 1. Figure
courtesy of C. R. Jackson, Global Ocean Associates.
Shanmugam 801
baroclinic flows occur when levels of constant
pressure are inclined to surfaces of constant den-
sity (Robertson and Ffield, 2005; Stewart, 2008). In
an oceanographic context, barotropic currents are
driven by the slope of the water surface, and these
currents are typical of the well-mixed shallower
(shelf) part of the ocean (Figure 2). In contrast,
baroclinic currents are driven by the vertical vari-
ations in the density of the ocean water caused by
changes in temperature and salinity. As a conse-
quence, baroclinic currents are commonly associ-
ated with internal waves and tides that propagate
along boundaries of density stratifications in the
deeper part of the ocean (Figure 2). Baroclinic cur-
rents can occur in mid-ocean depths (Figure 2)and
along the ocean floor of continental slopes and
submarine canyons. Despite its common usage in
oceanography, the baroclinic concept still remains
an unfamiliar theme in sedimentology.
The shelf edge is the defining bathymetric
boundary between the shallow mixed ocean and the
deep stratified ocean (Figure 2). The shelf-edge con-
cept is not applicable to a gently sloping carbonate
ramp setting or to periods of sea level lowstands.
According to the American Meteorological So-
ciety (Ocean Motion, 2012), a pycnocline is the
interface between the mixed and the deep ocean
layers where the density gradient is the greatest
(Figure 2). The density gradient is caused either by
differences in temperature (i.e., thermocline) or
by salinity (i.e., halocline). The ocean’s uppermost
100 m (328 ft) or so is well mixed by wind-driven
surface currents.
Internal waves are gravity waves that oscillate
along the interface between two water layers of
different densities, known as pycnoclines (Figure 2).
Although pycnoclines are primary boundaries of
density stratification for the existence of internal
waves, they are not essential in all cases. This is be-
cause any hydrostatically stable density stratification
is sufficient for sustaining internal waves (Garrett
and Munk, 1979). To distinguish these additional
boundaries from pycnoclines, the term “secondary
density stratification”is introduced in this article
(Figure 2). Internal waves are made visible at the
sea surface through the effect of internal wave cur-
rents on surface roughness (Gargett and Hughes,
1972). Internal waves are common phenomena in
Figure 2. Schematic diagram showing the position of the pycnocline (i.e., primary density stratification), where density gradient is the
sharpest, between mixed (upper) and deep (lower) ocean layers of different densities. Internal waves and tides propagate along the
boundaries of both primary and secondary density stratifications. Note that the shelf edge at 200 m (656 ft) is used as the defining
boundary between shallow-marine and deep-marine environments. Meteorological surface waves dominate shallow-marine (shelf)
environments, whereas oceanographic internal waves and astronomical internal tides propagate along the boundaries of density strati-
fication in deep-marine environments. Barotropic currents (red arrow) are generated by surface waves and tides, whereas baroclinic
currents (blue arrow) are generated by internal waves and tides. Note that baroclinic currents flow along density stratifications in open
water and along the sea floor. Relative increase in the density of fluid layer with increasing bathymetry is shown by r
1
,r
2
,r
3
,andr
4
. Note
that pycnoclines intersect only the sloping sea-floor topography but not the near-horizontal basin plain. The diagram is a composite
compilation of related concepts. This is partly based on Inman et al. (1976), Maxworthy (1979), Shanmugam (2008b), and Ocean Motion
(2012). Not to scale.
802 Internal Waves, Internal Tides, and Baroclinic Sands
coastal seas, open ocean, fjords, lakes, and the
atmosphere. However, the focus of this review is
on internal waves and tides in deep-marine en-
vironments (Figure 2).
Internal solitary waves, consisting of a single
isolated wave, are ubiquitous in stratified fluids.
Apel, (2002, p. 2) defined that “solitary waves are
a class of nonsinusoidal, nonlinear, more or less
isolated waves of complex shape, which occur com-
monly in nature. These waves maintain their co-
herence, and hence visibility, through nonlinear
hydrodynamics and appear as long, quasilinear
stripes in imagery.”Examples used in this article
are solitary type (i.e., solitons).
Internal tides are internal waves at a tidal fre-
quency (Shepard, 1975). The link between bottom
topography and the generation of internal (baro-
clinic) tides has been discussed by Robertson and
Ffield (2005). According to Garrett and Kunze
(2007, p. 57), “Internal tides are internal gravity
waves generated in stratified waters by the interac-
tion of barotropic tidal currents with variable bottom
topography.”
The term “tidalite”was originally introduced
for alternating units of traction and suspension de-
position from shallow-water tidal currents (Klein,
1971). The genetic term “internal tidalites”is pro-
posedinthisarticlefordepositsofinternaltidalcur-
rents. Pomar et al. (2012) introduced the genetic
term “internalite”for deposits of internal waves. This
is inconsistent with the conventional practice be-
cause the term “internalite”focuses on the position
internal, not the process wave (see Shanmugam,
2013). Deposits of baroclinic currents, associated
with both internal waves and tides, could be termed
“baroclinites.”
Tidal constituents refer to the influences of
the Earth’s rotation, the positions of the Moon
and the Sun relative to Earth, the Moon’s altitude
(elevation) above the Earth’s equator, and bathym-
etry of those that together cause tidal changes on
Earth. The equilibrium amplitude values of the
following tidal species are from Apel (1987).
Semidiurnal: n
1
=2
M
2
(period: 12.42 hr): principal lunar
(amplitude: 0.24 m [0.79 ft])
S
2
(period: 12.00 hr): principal solar
(amplitude: 0.11 m [0.36 ft])
N
2
(period: 12.66 hr): lunar elliptic
(amplitude: 0.05 m [0.16 ft])
K
2
(period: 11.97 hr): solar lunar
(amplitude: 0.03 m [0.10 ft])
Diurnal: n
1
=1
K
1
(period: 23.93 hr): solar lunar
(amplitude: 0.14 m [0.46 ft])
O
1
(period: 25.82 hr): principal lunar
(amplitude: 0.10 m [0.33 ft])
P
1
(period: 24.07 hr): principal solar
(amplitude: 0.05 m [0.16 ft])
Q
1
(period: 26.87 hr): elliptic lunar
(amplitude: 0.02 m [0.07 ft])
MODERN INTERNAL WAVES AND
INTERNAL TIDES
Origin
Internal waves are triggered by natural forces like
(1) wind (meteorological force), (2) tides (astro-
nomical force), (3) tropical cyclones (Nam et al.,
2007), (4) tsunamis (Santek and Winguth, 2007),
(5) river plumes (Nash and Moum, 2005), and by
man-made activities like sailing ships (Apel and
Gjessing, 1989). In this review, the focus is on nat-
ural processes.
Pycnoclines
Pycnoclines are the underpinnings of internal waves
and tides. Pycnoclines have been documented in a
wide range of water depths (2–2040 m [7–6691 ft])
in modern environments that include lakes, fjords,
bays, straits, submarine canyons and channels, car-
bonate banks, continental shelves, continental shelf
edges, continental slopes, and abyssal plains with
mid-ocean ridges and seamounts (Table 1). The
focus of this article is on oceanic pycnoclines that
intersect the slope at varying depths (Figure 2).
Pycnoclines are strongly controlled by both lati-
tude and season, and the temperature is far more
important than salinity in controlling vertical den-
sity profiles in most regions of the oceans of the
Shanmugam 803
Table 1. Bathymetry of Pycnoclines in Various Modern Environments (Lakes, Fjords, Bays, Straits, Submarine Canyons and Channels,
Carbonate Banks, Shelves, Shelf Edges, Slopes, and Abyssal Plains with Mid-Ocean Ridges)*
Example Environment Bathymetry of Pycnocline Reference
Location Number
in Figure 1
1 Scottish fjord: Gareloch 2–15 m (7–49 ft) Robins and Elliott (2009) –
2 Pacific Ocean: Monterey Bay 4 m (13 ft) Woodson et al. (2011) 3
3 Pacific Ocean: Mission Bay 5–10 m (16–33 ft) Lerczak et al. (2003) 5
4 Pacific Ocean: Sea of Okhotsk 10 m (33 ft) Sorokin and Sorokin
(1999, their figure 3)
41
5 Pacific Ocean: Kamchatka Shelf 14.5 m (48 ft) Serebryany and Pao (2008) 42
6 Atlantic Ocean: East Greenland Shelf 15 m (49 ft) Hirche and Bohrer (1987) 12
7 Pacific Ocean: Gulf of Alaska 19 m (62 ft) Walker et al. (1982; cited in
Churnside and Ostrovsky 2005)
1
8 Pacific Ocean: East China Sea,
Cheju Island
20 m (66 ft) Yanagi et al. (1996) –
9 Pacific Ocean: Yellow Sea 20 m (66 ft) Yanagi et al. (1996) 44
10 Pacific Ocean: South China Sea 20–60 m (66–197 ft) Wu et al. (2001, their figure 1) 45
11 Atlantic Ocean: Nigerian Shelf 20–35 m (66–115 ft) Amadi and Tobor (1988) 24
12 Pacific Ocean: Gulf of California 25–60 m (82–197 ft) Chen (1986) 6
13 North Sea 30 m (98 ft) Richardson et al. (2000) –
14 Atlantic Ocean: Bay of Biscay 40–70 m (131–230 ft) New and Pingree
(2000, their figure 16)
29
15 Atlantic Ocean: Caribbean Sea 40–100 m (131–328 ft) Martinez et al. (2007) 18
16 Atlantic Ocean: Florida-Atlantic Coast 45–60 m (148–197 ft) Leichter et al. (1998) 17
17 Atlantic Ocean: Iberian Peninsula 50 m (164 ft) da Silva et al. (2007) 30
18 Indian Ocean: Gulf of Oman 50 m (164 ft) Small and Martin (2002) 34
19 Indian Ocean: Andaman Sea 50 m (164 ft) Hyder et al. (2005) 40
20 Atlantic Ocean: Celtic Sea (shelf edge) 50 m (164 ft) Pingree and Mardell
(1985, their figure 9)
28
21 Indian Ocean: Mozambique Channel 60 m (197 ft) da Silva et al. (2009) 36
22 Indian Ocean: Krishna-Godavari Basin 60 m (197 ft) Ramana Murty et al. (2007) 39
23 Pacific Ocean: Monterey Canyon 60 m (197 ft) Hull et al. (2011) 4
24 Atlantic Ocean: New England Shelf 60 m (197 ft) Colosi et al. (2001) 14
25 Baltic Sea: Gotland Sea 60–80 m (197–262 ft) Wasmund et al. (1998) –
26 Atlantic Ocean: Gulf of Mexico 65–1000 m (213–3281 ft)
(varies with locations)
Vidal et al. (1994) Rubenstein
(1999) Herring (2010)
16
27 Atlantic Ocean: Portuguese Shelf
(shelf edge)
80–100 m (262–328 ft) Oliveira et al. (2002) 30
28 Atlantic Ocean: Middle Atlantic Bight
(Wilmington Canyon area) (shelf edge)
100 m (328 ft) Church et al. 1984) –
29 Atlantic Ocean: northwestern Africa
(shelf edge)
100 m (328 ft) Hagen (2001, his figure 3) –
30 Western Equatorial Atlantic: Northern
Brazil Basin
100 m (328 ft) Silva et al. (2009, their figure 2) 19
31 Pacific Ocean: Australian Northwest
Shelf (Scott Reef)
100 m (328 ft) Wolanski and Deleersnijder
(1998)
49
32 Pacific/Indian oceans: Lombok Strait 106 m (348 ft) Susanto et al. (2007, their
figure 4a)
48
33 Pacific Ocean: Oregon shelf 120 m (394 ft) (variable) Klymak and Moum (2003),
Moum et al. (2008)
2
804 Internal Waves, Internal Tides, and Baroclinic Sands
world (Pomar et al., 2012). Furthermore, the ba-
thymetry of pycnoclines is known to vary with time
at a given site (Martinez et al., 2007).
Physical Characteristics
Gill (1982) discussed the basic differences between
barotropic (surface) waves that develop at the air-
water interface and baroclinic (internal) waves that
develop at the water-water interface (Figure 3).
Fluid parcels in the entire water column move to-
gether in the same direction and with same velocity
in a surface wave, whereas fluid parcels in shallow
and deep layers of the water column move in op-
posite directions and with different velocities in an
internal wave (Figure 3). The surface displacement
and interface displacement are the same for a sur-
face wave, whereas the interface displacements are
large for internal waves. Although the free surface
movement associated with the baroclinic mode is
only 1/400 of the interface move movement, this is
still sufficient for baroclinic motions to be detect-
able by sea-surface changes (Wunsch and Gill, 1976).
The basic mechanism that initiates internal
gravity waves in the ocean is related to gravity and
buoyancy. The rate at which a stably stratified col-
umn of water oscillates under the combined influ-
ence of gravity and buoyancy forces is expressed
Table 1. Continued
Example Environment Bathymetry of Pycnocline Reference
Location Number
in Figure 1
34 Atlantic Ocean: Strait of Messina 80–130 m (262–427 ft) Casagrande et al. (2009) 32
35 Southeast Pacific Ocean 120–200 m (394–656 ft) Toniazzo et al. (2009) 20°S and
85° W
36 Pacific Ocean: Hawaiian Ridge 150–350 m (492–1148 ft) Karl and Lukas (1996);
Kang et al. (2000)
7
37 Atlantic Ocean: Middle Atlantic
(New York) Bight, shelf south of
Hudson Canyon (shelf edge)
200 m (656 ft) Gordon and Aikman (1981) 15
38 Atlantic Ocean: Great Meteor Seamount 200 m (656 ft) van Haren et al. (2004) 21
39 Pacific Ocean: Sulu Sea 200 m (656 ft) Liu et al. (1985) 46
40 Lake Baikal (eastern Siberia) 250 m (820 ft) Ravens et al. (2000) –
41 Indian Ocean: Red Sea 250 m (820 ft) Manasrah et al. (2006, their
figure 2)
33
42 Pacific Ocean: Sea of Japan-Korea Strait 300 m (984 ft) Gordon et al. (2002) 43
43 Antarctica: Western Weddell Sea 200–350 m (656–1148 ft) Gordon (1998, his figures 3a
and 3b)
51
44 Atlantic Ocean/Mediterranean:
Strait of Gibraltar
80 m (262 ft) (Mediterranean
side) 800 m (2625 ft)
(Atlantic side)
Alpers et al. (1996) 31
45 Northeastern Atlantic:
Faeroe-Shetland Channel
550–650 m (1804–2133 ft) Hall et al. (2011, their
figure 1b)
27
46 Atlantic Ocean: Mid-Atlantic Ridge 600 m (1969 ft) Apel (2002) –
47 Pacific/Indian oceans: Makassar Strait 680 m (2231 ft) Gordon (2005) 47
48 Northeastern Atlantic: Rockall and
Porcupine Banks and the Porcupine
Seabight (carbonate)
600–1000 m (1969–3281 ft) White and Dorschel (2010) –
49 Indian Ocean: Mascarene Ridge 1000–1200 m (3281–3937 ft) Morozov et al. (1999) 37
50 Atlantic Ocean: Rift Valley of the
Mid-Atlantic Ridge
2020–2040 m (6627–6693 ft) Thurnherr et al. (2002) –
*Most of these examples represent locations of case studies used in this paper (Figure 1; Table 2).
Shanmugam 805
by the Brunt-Väisälä frequency or buoyancy fre-
quency (e.g., Apel et al., 2006, their equation 10):
NðzÞ¼ ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi
g
r0
dr0
dz
s
where g= gravitational acceleration; r
0
= equili-
brium fluid density; z= height in fluid; and dr
0
/
dz= change in fluid density with height in fluid.
Internal solitary waves travel in packets. The
number of individual oscillations within the packet
increases as its age increases, with one new oscil-
lation added per Brunt-Väisälä period. The indi-
vidual oscillations are nonsinusoidal, with predom-
inantly downward displacements (Figure 4B). The
amplitudes are rank ordered, with the largest at
the front of the packet and the smallest at its rear
(Figure 4). The wavelengths and the crest lengths
are also rank ordered, with the longest waves at
the front of the group (Figure 5). Unlike surface
waves, internal waves can stretch over tens of kilo-
meters in length (Figure 5). Characteristically, a
younger (smaller) wave packet follows an older
(larger) packet forming a wave train in the Sulu
Sea (Figure 6). Unlike surface waves, internal waves
can propagate not only horizontally, but also ver-
tically and in any direction in between (Cacchione
and Pratson, 2004).
Figure 3. Diagram showing layer configuration in a two-layer fluid system for a barotropic (surface) wave (A) and a baroclinic (in-
ternal) wave (B) propagating from left to right (Gill, 1982). Also shown are the directions and velocities of flow at wave troughs and
crests. For the case shown, the lower layer is three times deeper than the upper layer and has 10% higher density. (A) Barotropic
(surface) wave showing the movement of fluid parcels in the entire water column (H) in the same direction (short horizontal arrows)
with same velocity at the crest. (B) Baroclinic (internal) wave showing the fluid-parcel movement in the upper (H
1
) and lower (H
2
) layers
in opposite directions (short horizontal arrows) with different velocities (U
1
and U
2
). (C) Equations of wavelength (l), wave number (k),
wave frequency (w), wave amplitude (h), and wave velocities (U
1
and U
2
). Tis the wave period (the interval of time during which a crest
travels one wavelength). Diagrams A and B are from Gill (1982, his figure 6.3), with permission from the Academic Press/Elsevier
Copyright Clearance Center’s RightsLink, License Number:3020870263403. The format of presentation is partly from MacKinnon (2012).
Additional equations and labels are inserted in this article for clarity.
806 Internal Waves, Internal Tides, and Baroclinic Sands
Apel (2002) summarized the physical proper-
ties of internal solitary waves. The characteristics
of internal solitary waves and surface waves are
summarized in Tables 2 and 3, respectively. Wave
amplitudes of internal solitary waves commonly
range from 5 to 50 m (16–164 ft), with a maximum
of 300 m (984 ft) (Table 2), whereas ampli-
tudes of surface waves are commonly less than 2 m
(6.56 ft) (Table 3). Internal waves on continental
shelvestendtobe10to30m(33–98 ft) in am-
plitude (Alford et al., 2012, their table 1). Internal
waves have their greatest wave height at inter-
mediate depths and their greatest velocities at the
bottom (LaFond, 1962). Although internal tides
have large amplitudes in the deep ocean, their sea-
surface height manifestations are only of a few
centimeters (Ray and Mitchum, 1997). This is
caused by the great increase in density difference
between air and water at the sea-surface interface in
comparison to the densitydifference betweenfluids
(i.e., water-water) at internal interfaces. For ex-
ample, the density of water is 1000 times greater than
that of air.
Wavelengths of internal solitary waves com-
monly range from 0.5 to 15 km (0.31–9 mi), with
a maximum of approximately 150 to 167 km (93–
104 mi) (Table 2), whereas those of surface waves
are less than 100 m (328 ft) (Table 3). Internal waves
commonly reveal longer wave periods (5–50 min)
than surface waves (9–10 s). Internal waves com-
monly travel at higher wave speeds (0.5–2ms
–1
[1.64–6.56 ft s
–1
]) than surface waves (25 cm s
–1
[10 in. s
–1
]) (Table 3). Like surface waves (Ray and
Mitchum, 1997), internal waves are spatially co-
herent and can propagate over great distances
(Inall et al., 2011).
Current Velocity
Deep-water bottom currents have been attributed
to internal waves in offshore California (Emery,
1956). Lonsdale et al. (1972) measured velocities
Figure 4. Schematic repre-
sentation of two individual pack-
ets of solitary internal waves
(nonlinear) in plan view (A) and
in profile view (B). The packet at
the right side represents the most
recent one generated by offshore
tidal flow at the shelf edge, and
the packet at the left side was
generated 12.5 hr earlier during
the previous semidiurnal tide.
The original diagram is from Apel
(2002) (from Jackson, 2004a,
with permission from Global
Ocean Associates. Additional
labels are inserted in this article
for clarity.
Shanmugam 807
of tidal currents as much as 17 cm s
–1
(7 in. s
–1
)over
the Horizon Guyot in the Mid-Pacific Mountains.
In this area, maximum semidiurnal current am-
plitudes reach 18 cm s
–1
(7 in. s
–1
) (Noble et al.,
1988). Instantaneous current speeds over the Cross
Seamount (290 km [180 mi] south of the island
of Oahu, Hawaii) exceeded 20 cm s
–1
(8 in. s
–1
)
(Noble and Mullineaux, 1989). Cacchione et al.
(1988) reported a maximum possible combined
flow speed of more than 30 cm s
–1
(12 in. s
–1
)from
the Horizon Guyot. Genin et al. (1989) measured
near-bottom currents for several days at three sites
on the summit of the Fieberling Guyot (32°26′N,
127°46′W). The observed currents were strong,
with maximum speeds of 48 cm s
–1
(19 in. s
–1
),
and diurnal currents were the dominant compo-
nent of the current field.
Brandt et al. (2002) reported results of high-
resolution velocity measurements conducted by
means of a vessel-mounted acoustic Doppler
current profiler on 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 pulselike
intense solitary disturbances propagated perpen-
dicular to the Brazilian Shelf, toward the north-
northeast. These internal waves were characterized
by maximum horizontal velocities of approximately
Figure 5. (A) Index map showing the study area in the Andaman Sea, Indian Ocean (arrowhead). Map credit: http://www.ngdc.noaa
.gov/mgg/topo/img/globeco3.gif (accessed October 1, 2012). (B) Satellite image showing both surface waves (near-horizontal faint
waves on the left side) and internal waves (long diagonal waves on the right side) in the Andaman Sea near the Andaman Islands (Indian
Ocean). Note that larger internal waves, with a wavelength of 5 km (3 mi), propagate in a westerly direction (arrow), whereas smaller
surface waves propagate in a north-south direction. Also note the largest separation of waves at the leading edge of the packet. Note the
active Barren Island Volcano emitting steam on the lower left side. This Advanced Land Imager (ALI) on the Earth Observing-1 (EO-1)
satellite acquired the image on March 6, 2007. The National Aeronautics and Space Administration (NASA) Earth Observatory image was
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 October 1, 2012).
808 Internal Waves, Internal Tides, and Baroclinic Sands
200 cm s
–1
(79 in. s
–1
) and maximum vertical ve-
locities of approximately 20 cm s
–1
(8 in. s
–1
).
Hyder et al. (2005) made observations of in-
ternal solitons that occurred between January and
April 1998 at a water depth of 440 m (1443 ft)
northeast of the Andaman Islands, Bay of Bengal.
Their observations indicated the occurrence of
internal solitons with thermocline depressions of
as much as 50 m (164 ft) and an upper-layer cur-
rent velocity of as much as 120 cm s
–1
(47 in. s
–1
).
These solitons only occurred during spring tides,
when the tidal range exceeded 1.5 m (5 ft).
Shepard et al. (1979), who measured current
velocities in submarine canyons, documented that
internal waves advance in both up- and down-
canyon directions. Measured values of velocity
reach as much as 100 cm s
–1
(39 in. s
–1
) in the up-
canyon direction and 265 cm s
–1
(104 in. s
–1
)inthe
down-canyon direction. Shepard (1975) suggested
that internal waves, which occur in canyon depths
Figure 6. (A) Index map showing the study area in the Sulu Sea (arrowhead). Map credit: http://www.ngdc.noaa.gov/mgg/topo/img
/globeco3.gif (accessed October 1, 2012). (B) Satellite image showing three packets (P1, P2, and P3) of internal waves, forming a wave
train (three yellow arrows) in the Sulu Sea between the Philippines (to the northeast) and Malaysia (to the southwest). This true-color Aqua
Moderate Resolution Imaging Spectroradiometer (MODIS) image was acquired on April 8, 2003. Image courtesy of Jacques Descloitres,
MODIS Land Rapid Response Team at National Aeronautics and Space Administration/Goddard Space Flight Center (GSFC). Image credit:
http://earthobservatory.nasa.gov/Newsroom/NewImages/images.php3?img_id=15334 (accessed October 1, 2012).
Shanmugam 809
Table 2. Empirical Data on Physical Characteristics of Internal Solitary Waves and Tides Used in This Study*
Location
Number
in Figure 1 Marine Setting Amplitude Wavelength
Wave
Period Wave Speed Bathymetry of Shelf Edges
1 Pacific Ocean: Gulf of Alaska (Churnside and
Ostrovsky, 2005; see also Walker et al., 1982)
2 m (6.56 ft) 4.6 km (2.86 mi) ––Shelf edge: 200 m (656 ft)
(Molnia, 1982)
2 Pacific Ocean: Oregon shelf (Moum et al.,
2003, 2007a)
40 m (131 ft) 0.17–0.22 km
(0.11–0.14 mi)
–0.6–0.8 m s
−1
(1.96–2.62 ft s
−1
)
Shelf edge: 200 m (656 ft)
(Moum et al., 2007b)
3 Pacific Ocean: Monterey Bay, California
(Cazenave, 2008; Woodson et al., 2011)
20–35 m (67–115 ft) 0.1–0.2 km
(0.06–0.12 mi)
–0.11 m s
−1
(0.36 ft s
−1
)
Shelf edge: 100 m (328 ft)
(Kunze et al., 2002, their figure 1)
4 Pacific Ocean: Monterey Canyon, California
(Broenkow and McKain, 1972; Petruncio
et al., 1998)
70–78 m (230–256 ft) 8.9–13.4 km
(5.53–8.33 mi)
–0.2–0.3 m s
−1
(0.65–0.98 ft s
−1
)
Shelf edge: 100 m (328 ft)
(Kunze et al., 2002, their figure 1)
5 Pacific Ocean: Mission Bay, California
(Lerczak et al., 2003)
–8 km (4.97 mi) –0.50 m s
−1
(1.64 ft s
−1
)
Shelf edge: 120 m (394 ft)
6 Pacific Ocean: Gulf of California (Apel
and Gonzalez, 1983; Fu and Holt, 1984;
Jackson, 2004a)
–0.2–1.6 km
(0.12–0.99 mi)
––Shelf edge depth not applicable
7 Pacific Ocean: Hawaiian Ridge (Ray and
Mitchum, 1997; Mitchum and Chiswell, 2000;
Kang et al., 2000; Rudnick et al., 2003)
20 m (66 ft) (common)
300 m (984 ft)
(maximum)
150 km (93 mi) –2.6 m s
−1
(8.53 ft s
−1
)
Shelf edge depth not applicable
8** Pacific Ocean: Horizon Guyot (Lonsdale
et al., 1972)
––––Ripples and dunes on Horizon
Guyot-top terraces; see location
7 for internal waves
9 Eastern Equatorial Pacific (Mack and
Hebert, 1999)
4 m (13 ft) 0.15–0.25 km
(0.09–0.16 mi)
––Shelf edge depth not applicable
10 Pacific Ocean: Galapagos Islands
(Jackson, 2004a)
–1 km (0.62 mi) –– –
11** Southeast Pacific, Southern Chile
(Sloyan et al., 2010)
––––Shelf edge: 200 m (656 ft)
(off southern Chile)
12 Atlantic Ocean: East Greenland shelf
(Jackson, 2004b)
–0.1 km (.06 mi) ––Shelf edge: 200–450 m (656–1476 ft)
(Dowdeswell et al., 1993); Shelf edge
at spill jet section: 250 m (820 ft)
(Brearley et al., 2012, their figure 3)
13 Atlantic Ocean: Nova Scotian shelf edge
(Smith and Sandström, 1988, their figure 8)
50 m (164 ft) –––Shelf edge: 230 m (755 ft)
(Smith and Sandström, 1988,
their figure 2B)
14 Atlantic Ocean: New England Shelf
(Colosi et al., 2001)
20–21 m (66–69 ft) 30 km (19 mi) –0.7 m s
−1
(2.29 ft s
−1
)
Shelf edge: 200 m (656 ft)
(USGS, 2007)
15 Atlantic Ocean: Mid-Atlantic (New York) Bight
(Apel et al., 1975; Jackson and Apel, 2002)
5–25 m (16–82 ft) 1–1.5 km
(0.62–0.93 mi)
8–25 min 0.5–1ms
−1
(1.64–3.28 ft s
−1
)
Shelf edge: 200 m (656 ft)
(Apel, 2002)
810 Internal Waves, Internal Tides, and Baroclinic Sands
16 Atlantic Ocean: Gulf of Mexico
(Rubenstein, 1999)
2–10 m (7–33 ft) 0.43 km (0.27 mi) –1.24 m s
−1
(4.06 ft s
−1
)
Shelf edge: 120–200 m (394–656 ft)
northwestern GOM (Slowey et al.,
2008)
Shelf edge: 50–120 m
(164–394 ft) northeastern GOM
(Koenig and Coleman, 2008)
17 Atlantic Ocean: Florida-Atlantic Coast
(Jackson and Apel, 2002)
–0.6 km (0.37 mi) ––Shelf edge: 70–100 m
(230–328 ft) (Reed, 2002)
18 Atlantic Ocean: Caribbean Sea
(Eakin et al., 2011)
30–50 m (98–164 ft) 0.3–3km
(0.19–1.86 mi)
––Shelf edge: 200 m (656 ft)
Puerto Rico (Schneidermann
et al., 1976)
19 Western equatorial Atlantic:
Northern Brazil Basin (Brandt et al., 2002; Ivanov
et al., 1993)
5–8 m (16–26 ft)
(common)
––2.0 m s
−1
(6.56 ft s
−1
)
Shelf edge: 100 m (328 ft)
Off the Amazon River (Milliman, 1979)
20 Western equatorial Atlantic:
Southern Brazil Basin (Pereira and Castro, 2007)
–5–100 km
(3.11–62.14 mi)
––Shelf edge: 80–130 m (262–427 ft)
Campos margin, (Viana et al., 1998)
21 Atlantic Ocean: Great Meteor Seamount
(Gerkema and van Haren, 2007)
Up to 75 m (246 ft) –––Shelf edge depth not applicable
22** Mid-Atlantic Ridge: Eastern Brazil Basin
(Polzin et al., 1997)
––––Shelf edge data not applicable
23 Atlantic Ocean: Moroccan shelf
(Vlasenko et al., 1996)
15–20 m (49–66 ft) 0.4–0.8 km
(0.25–0.50 mi)
–0.25–0.30 m s
−1
(0.82–0.98 ft s
−1
)
Shelf edge: 150 m (492 ft)
(Summerhayes et al., 1971)
24 Atlantic Ocean: Nigerian shelf (Jimoh, 2010) 8 m (26 ft) 50 km (31 mi) –0.53–0.76 m s
−1
(1.73–2.49 ft s
−1
)
Shelf edge: 100–130 m
(328–427 ft) (Akpati, 1983)
25 Atlantic Ocean: Southwest Africa
(Apel et al., 1975)
6–20 m (20–66 ft) 1.5–2.6 km
(0.93–1.62 mi)
12–55 min 0.5–1ms
−1
(1.64–3.28 ft s
−1
)
Shelf edge: ~200 m (656 ft)
(Dingle, 1973)
26 Arctic Ocean, Greenland Sea, and Barents Sea:
Spitsbergen Island (Kurkina and Talipova, 2011)
Up to 50 m (164 ft) 6–12 km
(3.73–7.46 mi)
–– –
27 Northeastern Atlantic Ocean: Faeroe-Shetland
Channel (Hall et al., 2011)
Up to 50 m (164 ft) 0.25–0.5 km
(0.16–0.31 mi)
–0.27 m s
−1
(0.88 ft s
−1
)
Shelf edge: 200 m (656 ft) West
Shetland shelf (Damuth and
Olson, 2001, their figure 1)
28 Atlantic Ocean: Celtic Sea (Pingree and
Mardell, 1981, 1985; Pingree et al., 1984;
Sharples et al., 2007)
50 m (164 ft) 30 km (19 mi) 30–40 min 0.70 m s
−1
(2.29 ft s
−1
)
Shelf edge: 200 m (656 ft)
(Pingree and Mardell, 1985)
29 Atlantic Ocean: Bay of Biscay (New and
Pingree, 1990, 2000)
25–30 m (82–98 ft) 1–3km
(0.62–1.86 mi)
30–40 min 1.1 m s
−1
(3.60 ft s
−1
)
Shelf edge: 200 m (656 ft)
(New and Pingree, 2000)
30 Atlantic Ocean: Iberian Peninsula (Apel 1979;
Small, 2002; da Silva et al., 2007)
50 m (164 ft) 2 km (1 mi) –0.56 m s
−1
(1.83 ft s
−1
)
Shelf edge: 200 m (656 ft)
(da Silva et al., 2007)
31 Atlantic Ocean/Mediterranean: Strait of Gibraltar
(Alpers et al., 1996; Lacombe and Richez,
1982; Farmer and Armi, 1988; Jackson and
Apel, 2002; Gómez-Enri et al., 2007)
50–80 m (164–262 ft) 0.6–0.8 km
(0.37–0.50 mi)
5–19 min 1–2ms
−1
(3.28–6.56 ft s
−1
)
Camarinal Sill is a critical factor
in propagating internal waves
(Figure 13E)
Shanmugam 811
32 Atlantic Ocean: Strait of Messina (Alpers and
Salusti, 1983; Casagrande et al., 2009).
40–50 m (131–164 ft) 1.3–2.5 km
(0.81–1.55 mi)
8–30 min 0.8–1ms
−1
(2.62–3.28 ft s
−1
)
Sill is a critical factor in propagating
internal waves (Figure 13D)
33 Indian Ocean: Red Sea (Gulf of Aqaba)
(Manasrah et al., 2006)
1.5 m (5 ft) 20 km (12 mi) –0.25–0.76 m s
−1
(0.82–2.49 ft s
−1
)
Shelf edge: 100 m (328 ft)
(Eilat subbasin)
Shelf edge: 70–80 m (230–262 ft)
(Aqaba subbasin) (Tibor et al., 2010)
34 Indian Ocean: Gulf of Oman (Small and
Martin, 2002)
25 m (82 ft) 41–42 km (25–26) –0.92 m s
−1
(3.01 ft s
−1
)
Shelf edge: 200 m (656 ft) (Small
and Martin, 2002, their figure 10)
35 Indian Ocean: Somalia, northeastern Africa
(Wang, 1997)
Up to 18 m (59 ft) 0.1–2.8 km
(0.06–1.74 mi)
29 min 0.87 m s
−1
(2.85 ft s
−1
)
Shelf edge: 200 m (656 ft) (Carbone
and Accordi, 2000); Figure 13C
36 Indian Ocean: Mozambique Channel
(da Silva et al., 2009)
–0.5–5km
(0.31–3.11 mi)
–1.35–1.55 m s
−1
(4.42–5.08 ft s
−1
)
Shelf edge: 200 m (656 ft)
(Lutjeharms, 2006)
37 Indian Ocean: Mascarene Ridge (Konyaev
et al., 1995; Morozov et al., 1999)
Up to 90 m (295 ft) 167 km (103 mi)
(estimated from
propagation distance)
–3ms
−1
(9.84 ft s
−1
)
Shelf edge depth not applicable
38 Indian Ocean: Arabian Sea (Murthy
et al., 1992)
40 m (131 ft) 45 km (28 mi) –0.83 m s
−1
(2.72 ft s
−1
)
Shelf edge: 100 m (328 ft)
(Murthy et al., 1992, their figure 1)
39 Indian Ocean: Bay of Bengal, Krishna-Godavari
Basin (Ramana Murty et al., 2007)
–2.1 km (1.30 mi)
(maximum)
––Shelf edge: 50 m (164 ft)
(Forsberg et al., 2007)
40 Indian Ocean: Andaman Sea (Osborne and
Burch, 1980; Jackson and Apel, 2002)
10–80 m (33–262 ft) 6–15 km (4–9 mi) 5–95 min 2 m s
−1
(6.56 ft s
−1
)
Shelf edge: 200 m (656 ft)
(Smith and Sandwell, 1997);
Figure 5
41 Pacific Ocean: Sea of Okhotsk (Nagovitsyn
et al., 1991; Mitnik and Dubina, 2007)
5–10 m (16–33 ft) 0.2–0.5 km
(0.12–0.31 mi)
(common) 5.8 km
(3.60 mi) (maximum)
10–15 min 0.6–1.4 m s
−1
(1.96–4.59 ft s
−1
)
Shelf edge: 200 m (656 ft)
(Gladyshev et al., 2003)
42 Pacific Ocean: Kamchatka shelf (Sabinin and
Serebryanyi, 2007, their figure 12)
Up to 17 m (56 ft) 0.4 km (0.25 mi) ––Shelf edge: 200 m (656 ft)
43 Pacific Ocean: Sea of Japan-Korea Strait
(Kim et al., 2001; Jackson, 2004c)
20–26 m (66–85 ft) 0.6–1.5 km
(0.37–0.93 mi)
3–10 min 0.5 m s
−1
(1.64 ft s
−1
)
Shelf edge: 100 m (328 ft)
East coast of Korea (Jackson,
2004c, his Figure 4)
44 Pacific Ocean: Yellow Sea (Hsu et al., 2000) 10 m (33 ft) 1 km (0.62 mi) ––Shelf edge: 100–200 m
(328–656 ft) (Isobe, 2008)
45 Pacific Ocean: South China Sea (Hsu et al.,
2000; Duda et al., 2004; Ko et al., 2008)
10 m (33 ft) (common)
200 m (656 ft)
(maximum)
1 km (0.62 mi) –1.45 m s
−1
(4.75 ft s
−1
)
Shelf edge: 200 m (656 ft)
(Xie et al., 2008)
Table 2. Continued
Location
Number
in Figure 1 Marine Setting Amplitude Wavelength
Wave
Period Wave Speed Bathymetry of Shelf Edges
812 Internal Waves, Internal Tides, and Baroclinic Sands
of as much as 3500 m (11,482 ft), were mostly
tidal in origin. Velocity measurements made by
Shepard et al. (1979) were not all baroclinic in ori-
gin. Shepard et al. (1979), who measured current
velocities commonly 3 m (10 ft) above the sea
bottom at varying ocean depths (46–4200 m [151–
13,779 ft]), did not claim that all their velocity
measurements were made along density stratifica-
tionsorpycnoclines.Nodocumentedexamplesof
pycnoclines that occur 3 m (10 ft) above the sea
bottom at a depth of, say, 4200 m (13,779 ft) in the
ocean exist. Most importantly, it is not practical to
distinguish barotropic tidal currents from baroclinic
tidal currents within submarine canyons during
velocity measurements. Mulder et al. (2012) be-
lieved that high-velocity currents in submarine can-
yons in the Bay of Biscay were related to internal
tides but did not provide empirical data in distin-
guishing baroclinic currents from barotropic cur-
rents. For this reason, the general term “deep-marine
tidal bottom currents”is preferred (Shanmugam,
2003).
In the Suruga Trough in Japan, semidiurnal
tidal fluctuations are evident in the current with
the total amplitude reaching 50 cm s
–1
(20 in. s
–1
)
at a depth of 1370 m (4494 ft). These currents have
been associated with internal tides (Matsuyama
et al., 1993). Velocity measurements associated
with internal tides in the Gaoping Submarine Can-
yon off southwestern Taiwan have revealed max-
imum velocities of more than 100 cm s
–1
(39 in. s
–1
)
(Lee et al., 2009). At these velocities, even gravel-
grade grains can be eroded, transported, and de-
posited by baroclinic tidal currents. Despite these
velocity measurements, observational data to link
types of sedimentary structures formed by internal
tidal currents with their velocities in modern deep-
marine environments are lacking.
Occurrence
Using mostly satellite synthetic aperture radar im-
ages, examples of modern internal waves and tides
in the oceans of the world were compiled by Apel
(2002), Jackson (2004a), and Jackson et al. (2012).
Most of the examples documented by Jackson
(2004a) represent coastal seas (Figure 1).
46 Pacific Ocean: Sulu Sea (Apel et al., 1985;
Jackson and Apel, 2002)
Up to 90 m (295 ft) 5–16 km (3–10 mi) 14–110 min 1.8–2.6 m s
−1
(5.90–8.53 ft s
−1
)
Shelf edge: 200 m (656 ft)
(Lewis, 1991); Figure 6
47 Pacific/Indian oceans: Makassar Strait
(Hatayama, 2004; Pujiana, et al., 2009)
100 m (328 ft) 10 km (6 mi) –1–1.5 m s
−1
(3.28–4.92 ft s
−1
)
Shelf edge: 185 m (607 ft)
(Prasetya et al., 2001, their figure 2)
48 Pacific/Indian oceans: Lombok Strait
(Susanto et al., 2005)
>100 m (328 ft) 2–5km(1–3mi) –1.97 m s
−1
(6.46 ft s
−1
)
Sill is a critical factor in propagating
internal waves (Figure 13F)
49 Pacific Ocean: Australian Northwest Shelf
(Holloway, 1994; Holloway et al., 1999)
25 m (82 ft) 20 km (12 mi) –0.4 m s
−1
(1.31 ft s
−1
)
Shelf edge: 100 m (328 ft) (Wolanski
and Deleersnijder, 1998)
50 Pacific Ocean: Poor Knights Islands
New Zealand (Stevens et al., 2005)
40 m (131 ft) 13 km (8 mi) 40 min 0.29 m s
−1
(0.95 ft s
−1
)
Shelf edge: 150–180 m
(492–591 ft) (Sim–Smith
and Kelly, 2008)
51 Antarctic Ocean (Southern Ocean): Weddell
Sea (Rees and Rottman, 1994; Robertson, 2005)
20 m (66 ft) 60–80 km
(37–50 mi)
–5.4 m s
−1
(17.71 ft s
−1
)
Shelf edge: 500–1000 m
(1640–3281 ft) (Barnes
and Lien, 1988)
*Bathymetric data of shelf edges are included for most cases. Satellite (synthetic aperture radar) images of most of these and other examples with basic references are available in two atlas volumes by the Global Ocean Associates
(Jackson and Apel 2002; Jackson 2004a). Depths of pycnoclines associated with internal waves in most of these locations are given in Table 1.
**Measurements of physical properties of internal waves are not available, but the oceanographic significance of internal waves has been discussed in the cited references.
Shanmugam 813
In a stratified ocean, internal solitary waves and
tides commonly originate above an area of steep
bathymetric variation, such as the shelf edge at a
depth of approximately 200 m (656 ft) (Apel et al.,
1975; Inman et al., 1976, their figure 4; Baines
1982; Pingree et al., 1986; Holloway, 1987; Apel,
2002, his figure 10; New and da Silva, 2002; Nash
et al., 2004; da Silva et al., 2007). Internal waves
that propagate toward onshore imply that wave
generation was at the shelf edge (Ostrovsky and
Table 3. Physical Characteristics of Different Types of Surface Waves that Include Surface Ocean Waves, Cyclonic Waves, Tropical
Instability Waves, Rogue Waves, and Tsunami Waves
Type of Wave
Wave Height (Amplitude =
1/2 the Wave Height)* Wavelength** Wave Period
†
Wave Speed
Surface ocean wave (Allen, 1970) <3.6 m (12 ft) (80% of
40,000 observations)
–– –
Surface ocean wave
(Komar, 1976; p. 78)
34 m (112 ft) (one of
the highest recorded)
–– –
Typical surface ocean wave
(Stewart, 2008)
–50–100 m
(164–328 ft)
––
Typical surface ocean wave
(Friedman and Sanders, 1978)
––~9 s ~25 cm s
−1
(10 in. s
−1
)
Surface cyclonic wave
(Hubbard, 1992)
––13–16 s (1989
Hurricane Hugo)
–
Surface cyclonic wave (Friedman
and Sanders, 1978)
–15–75 m
(49–246 ft)
––
Surface cyclonic wave
(Morton, 1988)
22 m (72 ft) (1969
Hurricane Camille)
–– –
Surface cyclonic wave
(Morton, 1988)
–––12 hours: shelf
duration
††
(1969
Hurricane Camille)
Surface swell (Shepard, 1973;
Friedman and Sanders, 1978)
15 m (49 ft) 300–900 m
(984–2953 ft)
6–14 s –
Tropical instability waves (Weisberg
and Weingartner, 1988)
–1100 km (3609 ft) –0.5 m s
−1
(1.6 ft. s
−1
)
Rogue waves, Draupner platform, North
Sea, January 1, 1995 (Haver, 2004)
26 m (85 ft) –– –
1958 Alaskan tsunami (Miller, 1960) 524 m (1719 ft)
(run-up)
–– –
The 1960 Chilean tsunami
(Takahasi and Hatori, 1961)
–500–800 km
(311–497 mi)
––
Typical Pacific tsunami waves
(Apel, 1987)
–––230 m s
−1
(755 ft. s
−1
)
2004 Indian Ocean tsunami wave
(NOAA, 2005; USGS, 2012;
Jaffe et al., 2005)
10–15 m (33–49 ft) ––222 m s
−1
(728 ft. s
−1
)
Common tsunami waves (Bryant, 2001) ––16–33 min –
Pacific tsunami waves (Apel, 1987) ––15–100 min (Pacific) –
*Wave amplitude = the vertical distance between the wave crest and the mean water level or one-half the wave height (sinusoidal); wave height = the vertical distance
between the wave crests and the wave troughs (sinusoidal); tsunami run-up height = the maximum height (reach) of tsunami wave landing with respect to mean sea
level (Shanmugam, 2006b).
**Wavelength = the horizontal distance between two adjacent wave crests or betwen two adjacent wave troughs (sinusoidal).
†
Wave period = the time required for one wavelength to pass a fixed point.
††
Shelf duration = the time a storm takes to cross the continental shelf (Morton, 1988).
814 Internal Waves, Internal Tides, and Baroclinic Sands
Stepanyants, 1989; Apel et al., 2006). Although in-
ternal waves and tides can propagate over continental
shelves and over gently sloping carbonate ramps,
such waves do not imply that they were generated
locally.
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 instead across the canyon and in
the direction parallel to the shelf break. Exam-
ples of this type are the (1) Hydrographer Can-
yon, U.S. Atlantic (Wunsch and Webb, 1979); (2)
Hudson Canyon, U.S. Atlantic (Hotchkiss and
Wunsch 1982); and (3) Monterey Canyon, U.S.
Pacific (Petruncio et al., 1998; Kunze et al., 2002).
In marine straits, such as the Strait of Gibraltar,
submarine ridges or sill are important topographic
features in generating internal waves (Gómez-Enri
et al., 2007).
Studies have examined the interaction of tidal
forces with submarine topography in understand-
ing ocean mixing. Polzin et al. (1997) have illu-
strated that interior mixing is concentrated over
the rough sea-floor topography of the Mid-Atlantic
Ridge in the Brazil Basin. This concept has been
illustrated by Jayne et al. (2004) for the Brazil Basin
(Figure 7). In this case, ubiquitous internal waves
are generated over the mid-ocean ridge by the
tides flowing over rough topography, whereas in-
ternal waves are conspicuously absent over the
smooth abyssal plains (Figure 7). Other examples
of ridges and seamounts are (1) the Mid-Atlantic
Ridge (15°W) and the Walvis Ridge (5°E) in the
Atlantic Ocean, (2) the southern edge of Mada-
gascar (45°E) and along the central Indian Ridge
(70°E) in the Indian Ocean, and (3) the South Fiji
Basin (170°E) and the seamounts of French Poly-
nesia (215°E) in the Pacific Ocean (St. Laurent
et al., 2002).
The advances in tidal mapping afforded by the
Topex-Poseidon satellite have allowed Egbert and
Ray (2000) to answer some longstanding questions
about tidal energy in the open ocean. In addressing
the tidal issues, Egbert and Ray (2000) focused on
the principal lunar semidiurnal M
2
tide because it
accounts for approximately two-thirds of the total
planetary dissipation (Cartwright and Ray 1991).
Analyses of the altimeter-derived cotidal charts re-
veal that, although most M
2
tidal energy dissipates
in shallow seas, approximately 1 TW (1 TW =
10
12
W) or 25 to 30% of the total energy dissipates
near rugged bottom topography, such as seamounts
and mid-ocean ridges in the deep open ocean. The
distribution of M
2
(semidiurnal) barotropic tidal
constituent in the oceans of the world clearly shows
that both shallow seas and deep oceans are affected
Figure 7. Schematic diagram showing
the generation of internal waves at a mid-
ocean ridge in the Brazil Basin in the South
Atlantic Ocean. Ubiquitous generation of
internal waves over the mid-ocean ridge is
caused by the tides flowing over rough
topography. In contrast, internal waves are
totally absent over the smooth abyssal
plains. Although most of the waves radiate
away from the ridge, some waves break up
into turbulence near the ridge and cause
turbulent mixing. In this article, it is sug-
gested that baroclinic currents associated
with internal waves rework sands on the
tops of ridges, seamounts, and guyots.
Additional labels are inserted in this article
for clarity. The original diagram is from
Jayne et al. (2004). The diagram is a
slightly modified version from St. Laurent
et al. (2012). Figure credit: The Oceano-
graphy Society.
Shanmugam 815
(Figure 8). In this semidiurnal distribution, note that
ocean tides generally rotate around amphidromic
points located offshore (Figure 8) and that flood and
ebb tides oscillate in opposite directions as the tidal
wave moves along.
A DEPOSITIONAL FRAMEWORK
Bathymetry of Shelf Edge and
Oceanic Pycnocline
The shelf edge is an important physiographic
boundary between the two major submarine prov-
inces, namely the continental shelf and the con-
tinental slope (Vanney and Stanley, 1983). Internal
solitary waves and tides commonly originate near
the shelf edge along oceanic pycnoclines. Kawabe
(1982, p. 115) states that, “In almost all coastal re-
gionsintheworldoceans,thepycnoclineinter-
sects the sloping bottom.”Empirical data show that
most pycnoclines (76%) intersect the sloping sea
floor in water depthsshallower than 200 m (656 ft)
(Figure 9). Pycnoclines intersect the sea floor
commonly near the shelf edge. For example, a
pycnocline intersects the shelf edge at the 100-m
(328-ft) isobath in the Wilmington Canyon area
of the Middle Atlantic Bight, (Church et al., 1984).
Figure 8. The M
2
(semidiurnal) barotropic tidal constituent in the oceans of the world, based on Topex-Poseidon satellite altimeter data.
The amplitude of the surface elevation in centimeters is indicated by the color bar scale. Note the high-amplitude locations (red-yellow)
in certain parts of the continental margins of the world. The white lines, which are cotidal differing by 1 hr, represent lines of constant
phase. Amphidromic points are spots where amplitude is zero and where phase lines merge. The curved arcs around the amphidromic
points show the direction of the tides, each indicating a synchronized 6-hr period. See Egbert and Ray (2000) for further details. Image
credit: Richard Ray, Space Geodesy branch, National Aeronautics and Space Administration–Goddard Space Flight Center; http://svs.gsfc
.nasa.gov/stories/topex/tides.html (accessed October 1, 2012).
816 Internal Waves, Internal Tides, and Baroclinic Sands
Similar instances of pycnoclines intersecting the
shelf edge at the 100-m (328-ft) isobath have been
reported from the northwestern African margin
(Hagen, 2001, his figure 3) and from the Portu-
guese Shelf (Oliveira et al., 2002). In the shelf area
south of Hudson Canyon in the Middle Atlantic
(New York) Bight, a pycnocline intersects the shelf
edge at 200 m (656 ft) (Gordon and Aikman,
1981). Off northwestern Australia, a pycnocline
intersects the sea floor near the shelf edge at ap-
proximately the 150- to 250-m (492–820-ft)
depths (Brink et al., 2007).
Empirical data show that most shelf edges (81%)
occur in water depths between 100 and 200 m (328
and 656 ft) (Figure 10). In the Antarctic conti-
nental margin, a shelf edge ranges between 500 and
1000 m (1640 and 3281 ft) in depth (Barnes and
Lien, 1988). In all these cases, the marked change in
sea floor topography associated with the shelf edge is
the primary factor that controls the generation and
propagation of internal waves and tides.
Continental Slopes and Submarine Canyons
At present, oceanographic models exist only for
explaining the origin of internal waves and tides
near the shelf edge (e.g., Hsu et al., 2000, their
figure 5; Sharples et al., 2007, their figure 1). These
Figure 9. (A) Conceptual diagram showing the intersection of pycnoclines with sloping sea-floor topography with increasing bathy-
metry. Note that most pycnoclines (76%) intersect the sea floor in water depths shallower than 200 m (656 ft). Pycnoclines that intersect
the sloping sea floor near the shelf edge (see text for examples) are of significance in this study. Hypothetical increase in the density of
fluid layer with increasing bathymetry is shown by r
1
,r
2
,r
3
,r
4
,andr
5
. (B) The number of cases plotted in different bathymetric
intervals is from Table 1. Note that Lake Baikal, a deep-water lake, is included in the percentage calculation (Table 1, number 40). (C)
Bathymetric intervals. These intervals are selected to be consistent with the intervals chosen for shelf edges (Figure 10).
Shanmugam 817
models were not designed to address the sedimen-
tologic aspects. Aspects of sediment movement
by internal waves and tides have been discussed by
various authors (Southard and Cacchione, 1972;
Cacchione and Southard, 1974; Cacchione et al.,
1988, 2002; McPhee-Shaw, 2006). To date, the
only convincing photographic documentation of
sedimentary bedforms, attributed to internal waves
and internal tides, in modern deep-marine envi-
ronments has been from seamounts and guyots in
the Pacific Ocean (Menard, 1952; Lonsdale et al.,
1972). These photographs are the only empirical
foundation for gaining insights into depositional
processes (e.g., traction vs. suspension) associated
with internal waves and internal tides. By integrating
available oceanographic and sedimentologic data, a
preliminary depositional framework is proposed
for continental slopes, submarine canyons, and guyots
(Figure 11).
In the vicinity of shelf-edge and associated
slopes, deposition is explained by three progres-
sive stages:
1. Incoming internal wave and tide stage: The in-
coming internal waves and tides, like tsunami
waves (Shanmugam, 2006a, b), simply repre-
sent transfer of energy. More importantly, these
incoming internal tides along pycnoclines are
sediment starved.
2. Shoaling transformation stage: As the internal
waves and tides encounter the sea floor near the
shelf edge, they tend to erode and incorporate
sediment. This sediment-entrainment process
commences once a significant frictional interac-
tion with the sea bottom exists. A similar trans-
formation occurs when waves of all kinds (i.e.,
sea-surface waves, storm waves, or tsunami waves)
approach the shoreface (Figure 11), which is
called “shoaling transformations”(Friedman and
Sanders, 1978).
3. Sediment transport and deposition stage: Fol-
lowing the sediment entrainment, the sediment-
charged flows transport sediment downslope (i.
e., return flows) and deposit them either on the
continental slope or in submarine canyons. Al-
though these return flows could broadly be
classified baroclinic in origin, they could behave
as mass-transport processes (e.g., debris flows),
turbidity currents, or bottom currents. The
nature and nomenclature of these outgoing
flows are yet to be established. Continental slopes
and submarine canyons are considered to be
Figure 10. Conceptual dia-
grams showing the frequency of
occurrences of shelf edges with
increasing bathymetry. (A)
Twelve percent of shelf edges
occur in water depths between
50 and 99 m (164 and 325 ft).
(B) Eighty-one percent of shelf
edges occur in depths between
100 and 200 m (328 and 656 ft).
(C) Five percent of shelf edges
occur in water depths between
201and500m(659and1640ft).
(D) Two percent of shelf edges
occur between 501 and 1000 m
(1644 and 3281 ft). Bathymetric
data are from Table 2.
818 Internal Waves, Internal Tides, and Baroclinic Sands
environments with high potential for deposi-
tion by these return flows or baroclinic events
(Figure 11).
Towering Guyots
Menard (1952, his plates 1 and 2) documented
symmetrical ripples using underwater photographs
from the Sylvania Seamount (present name: Wo-
dejebato Guyot, 11°53′N, 16°55′E) in the northern
Marshall Islands from a depth of 1372 m (4500 ft).
Importantly, he suggested the possibility that in-
ternal waves could have formed the ripples at these
great depths.
The Horizon Guyot (19°15′N, 160°00′W) is
a submarine volcanic ridge, 300 km (186 mi) long
and 75 km (47 mi) wide, located southwest of the
Hawaiian Ridge (Figure 1). It rises up 3558 m
(11,673 ft) from the surrounding abyssal sea floor
at a depth of 5000 m (16,404 ft) to the summit at
1442 m (4730 ft) (Figure 12). The truncated top
has been attributed to erosion during a previous
episode of subaerial exposure (Ladd and Newman,
1973). The flat top is composed of individual ter-
races that extend as much as 10 km (6 mi) in length
and 3 km (1.86 mi) in width. Lonsdale et al. (1972)
documented symmetrical ripples, asymmetrical rip-
ples (Figure 13), and asymmetrical dunes (Figure 13)
in foraminiferal sands on the terraces. Lonsdale et al.
(1972) suggested that internal tidal currents were
responsible for forming these bedforms. Although
cores from these specific bedforms are lacking for
confirming their true depositional origin, a rework-
ing origin by internal tidal currents is a good pos-
sibility for three reasons: (1) foraminiferal sands
were available for reworking on these terraces, (2)
flat terrace surfaces were suitable for reworking by
the internal tidal currents, and (3) current meters,
deployedwithin12m(39ft)ofthebottom,mea-
sured strong tidal currents (see “Current Velocity”
section above). This scenario may allow for the re-
working of foraminiferal sands by baroclinic currents
on the guyot terraces (Figure 12).
The Great Meteor Seamount (30°00′N, 28°
30′W) in the Atlantic Ocean (Figure 1) is another
example that had favorable conditions for current
reworking. (1) It rises up from a depth of 4800 m
(15,748 ft) to approximately 270 m (886 ft) under
the surface of the sea (Gad and Schminke, 2004).
(2) A pycnocline has been reported at a depth of
200 m (656 ft) (van Haren et al., 2004). (3) Internal
tides with large amplitudes of as much as 75 m
Figure 11. Conceptual oceanographic and sedimentologic framework showing deposition from baroclinic currents on continental
slopes, in submarine canyons, and on guyots. On continental slopes and in submarine canyons, deposition occurs in three progressive
stages: (1) incoming internal wave and tide stage, (2) shoaling transformation stage, and (3) sediment transport and deposition stage.
Continental slopes and submarine canyons are considered to be environments with high potential for deposition from baroclinic
currents. In the open ocean, baroclinic currents can rework sediments on flat tops of towering guyot terraces, without the need for the
three stages required for the deposition on continental slopes. In this model, basin plains are considered unsuitable environments for the
deposition of baroclinic sands. Not to scale.
Shanmugam 819
(246 ft) have been reported (Gerkema and van
Haren, 2007). (4) The top surface of the seamount
is covered by coarse biogenic sand composed of
fragments of corals and of mollusks as the result of
erosion of older coral reefs, which fringed the top
of the seamount when the sea level was lower
(Nellen, 1998). (5) The carbonate sands are rip-
pled and scoured on the top of seamounts, and the
sands are commonly rather coarse because the
finer fractions are removed by the action of strong
currents (Levin and Nittrouer, 1987).
Unlike continental slopes and submarine can-
yons that require three stages for deposition asso-
ciated with internal waves and tides, the reworking
of sediment on the flat tops of guyots commences
once the incoming internal tidal currents interact
with the sediment.
Basin Plains
Modern examples where pycnoclines with internal
tides directly intersect the abyssal sea floor do not
exist. This is because abyssal sea floors do not con-
stitute a sloping topography (Figure 2). Empirical
data have shown the absence of internal waves in
the abyssal plains (Figure 7). Therefore, basin plains
have only low or no potential for sand deposition by
baroclinic currents (Figure 12).
SEDIMENTOLOGIC CHALLENGES
Process sedimentology of internal waves and tides
is an unknown entity. Therefore, sedimentologic
challenges associated with interpreting ancient deep-
marine sands of baroclinic origin are innumerable.
Figure 12. An air-gun seismic profile showing the towering
nature of the Horizon Guyot with three terraces, Mid-Pacific
Mountains. A hypothetical position of baroclinic tidal currents is
added in this article for illustrating their mid-ocean position with
respect to terraces. The potential for reworking by baroclinic tidal
currents is moderate for guyot terraces, whereas the reworking
potential is absent or low for the surrounding abyssal floor (see
Figure 7). The guyot summit occurs at a depth of 1442 m (4731 ft)
(Lonsdale et al., 1972, their figure 2B). The seismic profile is across
the northern slope of the guyot. Credit: Lonsdale et al. (1972, their
figure 13, left profile), with permission from the Geological Society
of America.
Figure 13. Cross profile showing asymmetrical dunes and asymmetrical ripples observed from side-looking sonar and photographic
evidence obtained from the terrace of the Horizon Guyot, Mid-Pacific Mountains. Bathymetry of bedforms: 1630 to 1632 m (5346–5353 ft).
Dune heights (H) were estimated from the length of acoustic shadows. Redrawn from Lonsdale et al. (1972, their figure 10), with per-
mission from the Geological Society of America. b= dip angle.
820 Internal Waves, Internal Tides, and Baroclinic Sands
Some of these issues have recently been discussed in
response to the local outcrop studies in China and
Spain (Shanmugam, 2012a; 2013). However, the
following narrative provides a broader global per-
spective on 12 critical points.
Lack of Core Data from Modern Settings
In a pioneering article on process sedimentology,
Sanders (1963, p. 178) emphasized the importance
of primary sedimentary structures that yield in-
sights into the fluid mechanics of moving currents
in terms of the fundamental differences between
the mechanisms of suspension and traction. Klein
(1975) has used the physical sedimentary struc-
tures as evidence for the alternating traction and
suspension phases of deposition in support of his
interpretation of sediments in Deep Sea Drilling
Project Leg 30 cores from the western equatorial
Pacific as deposits of deep-water tidal bottom cur-
rents (i.e., tidalites). Other researchers who have
interpreted sediments and sedimentary rocks as
deep-water tidalites using sedimentary structures
exist (Laird, 1972; Skilbeck, 1982; May et al.,
1983; Galloway and Hobday, 1983; Mutti, 1992;
Cowan et al., 1998; Shanmugam, 1997, 2002, 2003;
Shanmugam et al., 2009; Mutti and Carminatti,
2012).
The advantage of studying modern sediments
is that it allows one to make a direct link between
a depositional process and its product. In the case
of modern subtidal settings, for example, Visser
(1980) established the link between double mud
layers and alternating ebb and flood tidal currents
with extreme time-velocity asymmetry. Such ob-
servations from modern settings are totally lacking
for deposits of internal waves and internal tides.
Note the following examples.
1. As pointed out previously, Menard (1952, his
plates 1 and 2) documented symmetrical ripples
using underwater photographs from the Sylva-
nia Seamount in the northern Marshall Islands
and attributed their origin to internal waves.
However, no sediment cores were taken through
these bedforms to establish their internal sedi-
mentary structures and their true origin.
2. Stride and Tucker (1960) attributed the devel-
opment of modern sand waves near the shelf
edge to internal waves, again, without cores.
3. Lonsdale et al. (1972) documented modern trac-
tionbedforms,suchasasymmetrical,ripples,and
asymmetrical dunes (Figure 13)ontheHorizon
Guyot, Mid-Pacific Mountains, using sonar and
bottom photographic images, and interpreted
these bedforms as products of reworking by in-
ternal tidal currents. However, the limitations of
this study were twofold. First, dune heights were
estimated from the length of acoustic shadows.
Second, no sediment cores were taken directly
from these traction bedforms. Without sediment
cores through the bedforms, it is impossible
to determine the types of internal sedimentary
structures associated with these bedforms. There-
fore, tidal features, such as alternating traction
and suspension units, mud drapes, double mud
layers, and others, cannot be confirmed.
4. Karl et al. (1986), using sparker profiles, docu-
mented sand waves in the heads of the sub-
marine 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 from these sand waves. Karl et al. (1986)
speculated that internal waves were responsible
for the origin of sand waves.
5. Quaresma et al. (2007) studied nonlinear in-
ternal waves propagating over the northern shelf
of Portugal, near the Nazaré submarine canyon.
Although sediment samples with 93% sand are
available from this area, no descriptions of pri-
mary sedimentary structures in sands exist, nor
do any discussions of depositional mechanisms
of sand by baroclinic currents.
6. Martini et al. (2011) studied shoaling of internal
tides on the Oregon continental slope during
periods of spring and neap tides. However, as
oceanographers, their objective was not to study
sedimentary structures using sediment cores.
In summary, core data on sandy deposits of in-
ternal waves and tides with sedimentary structures
in modern deep-marine settings are totally absent.
Shanmugam 821
This is the single most important challenge to se-
dimentologists because, without the direct link
between primary sedimentary structures and de-
positional mechanics of baroclinic currents, there
is no process-sedimentologic basis for interpreting
ancient baroclinic sands.
Figure 14. Maps showing the variable directions of propagation of internal waves with respect to shoreline or shelf edge seen as
surface manifestations on satellite images. (A) Internal waves propagating toward the shoreline of Palawan Island in the Sulu Sea
(Figure 6). The shoreward propagation of internal waves was also reported from the South China Sea (Hsu et al., 2000, their figure 2).
(B) Internal waves propagating away from the shoreline or shelf edge in the Yellow Sea (Hsu et al., 2000, their figure 8); similar
examples have been reported from the Bay of Biscay (New and da Silva, 2002) and from the Iberian Peninsula (da Silva et al., 2007).
(C) Internal waves propagating nearly parallel to the shoreline of northern Somalia in the Indian Ocean (Jackson, 2004d, his figure 3).
In southern California, off Mission Bay, internal tides on the slope and shelf break were dominated by alongshore-propagating coastal
trapped waves (Lerczak et al., 2003). Off East Greenland, internal solitary waves were propagating northward parallel to the coast on
May 14, 2001, at 1350 coordinated universal time (UTC) (Jackson, 2004b, his figure 2). (D) Internal waves propagating parallel to the
strait or channel axis in the Strait of Messina (based on the satellite image from the National Aeronautics and Space Administration
(NASA, 2012), Jet Propulsion Laboratory, California Institute of Technology, NASA/The Goddard Space Flight Center (GSFC), Ministry of
Economy, Trade, and Industry (METI), Government of Japan Earth Remote Sensing Data Analysis Center (ERSDAC), The Advanced
Spaceborne Thermal Emission and Reflection Radiometer (ASTER), Japanese Resource Observation System (JAROS), and U.S./Japan
ASTER Science Team; http://asterweb.jpl.nasa.gov/gallery-detail.asp?name=messina-wave (accessed October 1, 2012). Note the sill at
the point of origin of internal waves (Brandt et al., 1997). (E) Internal waves propagating in the same direction on both sides of the
Strait of Gibraltar. Note the position of the Camarinal Sill at the point of origin of internal waves (Gómez-Enri et al., 2007). The sill has
been attributed to two different origins: (1) megaslide (Blanc, 2002) and (2) tectonic crustal shortening (Luján et al., 2011). (F) Internal
waves propagating in opposite directions from the point of origin, which is a sill in the Lombok Strait (Susanto et al., 2005). The features
shown are schematic and not to scale.
822 Internal Waves, Internal Tides, and Baroclinic Sands
Paleocurrent Directions
In distinguishing deposits of internal waves and
internal tides in the ancient stratigraphic record,
bidirectional cross-bedding has been used as an im-
portant criterion (Gao and Eriksson, 1991; He et al.,
2008). This is based on the notion that up- and
down-currents in channel environments develop
bidirectional cross-bedding. However, satellite im-
ages of modern internal waves reveal that the direc-
tions of propagation of internal waves are highly
variable with respect to the shoreline, the shelf edge,
and the channel axis (Figure 14). Selected examples
are (1) internal waves that propagate toward the
shoreline; (2) internal waves that propagate away
from the shoreline or the shelf edge; (3) internal
waves that propagate nearly parallel to the shore-
line; (4) internal waves that propagate in the direc-
tion parallel to the strait axis or channel axis, con-
trolled by a sill; (5) internal waves that propagate in
the same direction on both sides of a strait, con-
trolled by a sill; and (6) two wave trains of internal
waves propagating in opposite directions from the
point of origin, a sill, in a strait (Figure 14).
However, no systematic linking exists of wave-
propagation directions seen as the sea-surface man-
ifestations on satellite images (Figures 5,6)with
their respective influence on internal sedimentary
structures (i.e., dip directions) in the depositional
bedforms on the modern sea floor. This lack of a
link between the direction of wave propagation
along pycnoclines and the direction of current
movement on the sea floor is further compounded
by the presence of local sills on the sea floor, which
invariably control the direction of wave propagation
(Figure 14D, E, F). Furthermore, Dykstra (2012, his
figure 14.3B caption) states that, “If more than one
wave is present in the ocean at different depths, which
can occur in well-stratified water with significant sea-
floor topography (e.g., Robertson, 2005), current di-
rections along the sea floor can become quite com-
plicated.”The other complication is that, in some
submarine canyons, tidal currents flow across the
canyon and in the direction parallel to the shelf
break (Allen and Durrieu de Madron, 2009). Im-
portantly, the development of bidirectional cross-
bedding in modern submarine canyons or channels has
never been documented using sediment core. Under
this umbrella of knowledge vacuity, any interpretation
of paleocurrent directions as evidence for deposition
by baroclinic currents in outcrop studies is sedi-
mentologically erroneous.
Vertical Facies Models
Gao and Eriksson (1991, their figure 3B), strictly
from the study of outcrops of the Middle Ordovi-
cian in the southern Appalachians, first developed
three vertical facies models for internal-tide depos-
its in submarine channel environments (Figure 15A,
B, C). He et al. (2008, their figure 2) and He et al.
(2011, their figure 11) have proposed a total of
seven vertical facies models for internal-tide de-
posits and internal-wave deposits. These models
represent (1) three submarine channel models of
Gao and Eriksson (1991), (2) three unchannelized
continental slope models (He et al., 2008, their
figure 2D, E, F), and (3) an abyssal basin model
(Figure 15D). However, empirical data show that
abyssal basins are unlikely environments for sand
deposition by baroclinic currents (Figure 7). The
other problem is that a particular vertical trend
can be present in more than one environment. For
example, a key vertical trend of internal-tide de-
posits of submarine continental environments is
the upward-coarsening trend with bidirectional
cross-bedding (Figure 15A). However, such trends
are also common in deposits unrelated to internal
waves and internal tides in channelized environ-
ments. For example, upward-coarsening trends with
bidirectional cross-bedding have been documented
in estuarine tidal sand bars (Shanmugam et al.,
2000, their figure 9). At present, criteria to dis-
tinguish deposits of barotropic tidal currents in
shallow-marine environments from those of baro-
clinic tidal currents in deep-marine environments
are not available. Furthermore, upward-coarsening
trends are considered typical of storm deposits
(Bádenas and Aurell, 2001; Pomar et al., 2012).
The uncertainty of outcrop-based vertical facies
models has long been recognized for storm (Dott
and Bourgeois, 1982), fluvial (Miall, 1985), and
turbidite (Shanmugam et al., 1985) deposits. Even
the classic Bouma sequence is considered obsolete
Shanmugam 823
because of the lack of theoretical, experimental,
and empirical bases (Shanmugam, 2012c).
Internal-Wave Deposits Versus
Internal-Tide Deposits
Internal waves can be distinguished from internal
tidesinmodernoceansbymonitoringtidalfre-
quency. However, the distinction between deposits
associated with internal waves and those associated
with internal tides in the ancient stratigraphic rec-
ord has never been resolved using core studies of
modern analogs. Perhaps, because of this knowl-
edge gap, He et al. (2008, 2011) have combined
characteristic structures of both internal-wave and
internal-tide deposits. These sedimentary structures
are (1) bidirectional cross-lamination, (2) cross-
lamination dipping upslope, (3) multidirectional cross-
lamination, (4) flaser bedding, (5) wavy bedding,
(6) lenticular bedding, (7) double mud layers, and
(8) reactivation surfaces. The problem is that these
sedimentary structures are associated with de-
posits of tidal currents not only in shallow-marine
environments (Reineck and Wunderlich, 1968;
Klein, 1970; Visser, 1980; Terwindt, 1981; Allen,
1982; Nio and Yang, 1991; Dalrymple, 1992;
Alexander et al., 1998; Archer, 1998; Shanmugam
et al., 2000; Davis and Dalrymple, 2012), but also
in deep-marine environments (Klein, 1975; Cowan
et al., 1998; Shanmugam, 2003; Shanmugam et al.
2009; Mutti and Carminatti, 2012). More impor-
tantly, none of the structures listed above can be
used reliably to distinguish internal-wave depos-
its from internal-tide deposits. The current practice
of bundling together internal-wave deposits with
internal-tide deposits under a composite category
defies the very purpose of process sedimentology
(Sanders, 1963).
Turbidites
Turbidity current is a sediment-gravity flow with
Newtonian rheology and turbulent state in which
sediment is supported by fluid turbulence and
Figure 15. Vertical facies
models of internal-tide deposits
in submarine channel (A, B, C)
and internal-tide and internal-
wave deposits in abyssal basin
(D) environments. In facies D,
Ta, Tc, and Te notations of the
Bouma sequence are added in
this article. Credit: Facies models
A, B, and C are from Gao and
Eriksson (1991), with permission
from the Geological Society of
America; Facies model D is from
He et al. (2011), with permission
from Springer. Copyright Clear-
ance Center’s RightsLink, License
Number: 995890769723.
824 Internal Waves, Internal Tides, and Baroclinic Sands
from which deposition occurs through suspension
settling (Dott, 1963; Sanders, 1965; Middleton
and Hampton, 1973; Shanmugam, 1996, 2012c;
Talling et al., 2007). Deposition from turbidity
currents occurs when the velocity of turbidity cur-
rents decreases with time (waning flow), resulting
in normal grading (Dott, 1963; Sanders, 1965;
Shanmugam. 1996). Turbidites commonly exhibit
erosional basal contacts (Bouma, 1962; Middleton
and Hampton, 1973; Mutti, 1992). The importance
of turbulence in internal waves and the possibi-
lity of forming normal grading have been discussed
by Kneller et al. (1997). Various features, such as
(1) exhumed terraces, (2) truncation of chert lay-
ers, (3) erosion of pelagic sediment cap, (4) ero-
sional steepening, (5) sudden thinning of sediment
by erosional beveling, and (6) scour marks on the
Horizon Guyot in the Mid-Pacific Mountains, were
attributed to erosion by internal tidal currents
(Lonsdale et al., 1972). If so, baroclinic sands in the
rock record may exhibit normal grading with ero-
sional basal contacts, not unlike turbidites in deep-
marine environments and storm deposits in shal-
low-marine environments (Kreisa, 1981). Given
the current lack of sedimentologic criteria for dis-
tinguishing baroclinic sands, interpretation of
deep-water sands as turbidites should proceed with
caution.
Mass-Transport Deposits
The significance of sandy mass-transport deposits
in petroleum exploration has recently been re-
viewed (Shanmugam, 2012c). In a study of the
Horizon Guyot based on textural trends and pres-
ervation of bedforms on the pelagic cap, Kayen et al.
(1989) suggested that current-generated sediment
transport direction was upslope. Kayen et al. (1989)
also reported that slumping of the sediment cap
occurred on the northwestern side of the guyot on a
1.6° to 2.0° slope in the zone of enhanced current
activity. These slump blocks appear to be discrete
and to have a relief of 6 to 15 m (20–49 ft), with
nodular chert beds cropping out along the head-
wall of individual rotated blocks. The implication is
that ancient baroclinic deposits may be composed
of interbedded slump and internal tidalite facies.
Similar facies association of mass-transport deposits
and tidalites has been reported from submarine can-
yon settings (May et al., 1983; Shanmugam, 2003).
Bottom-Current Reworked Sands
The significance of bottom-current reworked sands
in petroleum exploration has recently been reviewed
(Shanmugam, 2012c). Southard and Stanley (1976)
recognized five types of bottom currents at the shelf
edge. These currents are generated by (1) thermo-
haline differences, (2) wind forces, (3) tidal forces,
(4) surface waves, and (5) internal waves. In this re-
view, four types of deep-water bottom currents,
namely (1) thermohaline-induced geostrophic bot-
tom currents (i.e., contour currents), (2) wind-driven
bottom currents, (3) deep-marine tidal bottom cur-
rents, and (4) baroclinic currents, are considered
(Shanmugam, 2008b).
Traction structures are considered to be an in-
tegral part of contourites (i.e., deposits of contour
currents) (Hubert, 1964; Hollister, 1967; Mutti, 1992;
Shanmugam, 2000, 2008b; Ito, 2002; Martýn-
Chivelet et al., 2008; Mutti and Carminatti, 2012).
Similarly, traction structures have been attributed
to bottom-current reworking by the wind-driven
Loop Current (Shanmugam et al., 1993a, b). These
deposits are characterized by cross-bedding, ripple
lamination, and horizontal lamination (Figure 16). It
would be a challenge to distinguish these bottom-
current reworked sands with traction structures from
baroclinic sands that also develop traction bedforms
(Figure 13).
Storm-Related Deposits
Although storms are compared to internal waves
(Pomar et al., 2012), major differences and simi-
larities are observed between the two. For example,
(1) the Atlantic Oceanographic and Meteorological
Laboratory (AOML, 2011) defines a tropical storm
as a type of tropical cyclone (broader category) with
amaximumsustainedwindvelocityof62to119km
hr
–1
(39–74 mi hr
–1
). However, internal waves are
not defined on the basis of velocity. (2) The “storm”
phenomenon represents multiple depositional
processes (Johnson and Baldwin, 1996, p. 249),
Shanmugam 825
whereas baroclinic tidal currents represent a single
process. (3) Hummocky cross-stratification has
been attributed to both storms (Harms et al., 1975)
and internal waves (Pomar et al., 2012). (4) Up-
ward-coarsening trends have been related to both
storm deposits (Bádenas and Aurell, 2001; Pomar
et al., 2012), and internal-wave and internal-tide
deposits (Figure 15A). Therefore, vertical facies
trends are unreliable templates for distinguishing
baroclinic sands.
Tsunami-Related Deposits
Despite an influx of publications on criteria for
recognizing tsunami deposits, the challenge of dis-
tinguishing tsunami-related deposits in the rock
record remains (Shanmugam, 2006b, 2012b). The
notion that a tsunami represents a single (unique)
depositional process is a myth because a tsunami is
a triggering event of multiple depositional processes
(Shanmugam, 2012b). In the ancient geologic record,
Figure 16. Summary of traction features in-
terpreted as indicative of deep-water bottom-
current reworking (from Shanmugam et al.,
1993a, with permission from AAPG).
826 Internal Waves, Internal Tides, and Baroclinic Sands
distinguishing internal tidalites from tsunami deposits
poses an even greater challenge because neither de-
posit has an established set of criteria for recognition.
Evidence for Pycnoclines
The supreme evidence for interpreting internal ti-
dalites in the rock record is the physical evidence
for pycnoclines (Shanmugam, 2013). Without that
evidence for density stratification, no difference be-
tween a surface tidalite formed by surface (baro-
tropic) tides on a shallow-marine shelf and an in-
ternal tidalite formed by internal (baroclinic) tides
in a deep-marine slope or canyon environment ex-
ists. The interpretations of ancient strata as de-
posits of internal waves and tides by He et al.
(2011) were not based on the ultimate evidence
for internal-wave deposits, which is the pycno-
cline (Shanmugam, 2012a). In defending their
interpretations, He et al. (2012, p. 371) state,
“Conclusive evidence for the existence of a pycno-
cline in our stratigraphic record is currently lacking.
Because the absence of proof is not proof of the
contrary, it is unreasonable to use this as a basis
to negate the possibility that these deposits may
have been generated by internal waves and internal
tides.”In discussing the criteria for recognizing
internal-tide deposits, Dykstra (2012) has over-
looked the importance of paleopycnoclines.
Biological and sedimentologic criteria for rec-
ognizing paleopycnoclines in the rock record
have been discussed by various authors (e.g.,
Byers, 1977; Woodrow, 1983; Ettensohn and Elam,
1985). In developing criteria for paleopycno-
clines, Woodrow (1983) suggested that the rock
units deposited in water depths shallower than the
pycnocline should be characterized by deposits of
wind-driven bottom currents, whereas the units
deposited in water depths deeper than the pycno-
cline should be dominated by turbidites. These
criteria are flawed for the following key reasons:
1. Pycnocline depths in the modern Gulf of Mex-
ico are highly variable depending on locations.
The depth ranges from 65 m (213 ft) (Rubenstein,
1999) to 1000 m (3281 ft) (Herring, 2010). Nev-
ertheless, wind-driven bottom currents have
been documented to develop ripples at 3091 m
(10,141 ft) of water depth on the modern sea
floor in the Gulf of Mexico (Pequegnat, 1972).
Clearly, these ripples occur in depths that are
much deeper than the pycnocline depth, defy-
ing the criterion.
2. It has been documented that pycnoclines can
develop at great depths of more than 2000 m
(6562 ft) (Table 1). Also well known is that
turbidites can occur at water depths shallower
than the pycnoclines in a wide range of envi-
ronments (Shanmugam, 2006a). Thus, the tur-
bidite criterion is flawed.
Until we develop objective criteria for recog-
nizing pycnoclines in the rock record, interpre-
tations of baroclinic sands will always remain
dubious.
Seismic Wave Geometry
Reeder et al. (2011, their figure 4) reported large
subaqueous sand dunes with seismic wave geom-
etry on the upper continental slope of the northern
South China Sea. These dunes, composed of fine
to medium sand, occur in water depths of 160 to
600 m (525–1969 ft). Their amplitudes and wa-
velengths are greater than 16 and 350 m (52 and
1148 ft), respectively. Reeder et al. (2011) suggest
that these sand dunes are formed by the largest ob-
served internal solitary waves of the world by tidal
forcing with amplitudes exceeding 100 m (328 ft).
Although internal tides might conceivably develop
large sand waves on the continental slope, uncer-
tainties associated with the data and the interpreta-
tion of Reeder et al. (2011) exist. Note the following
problem areas.
1. The reported sand-size grade is based on four
surficial sediment samples. These samples do not
reveal anything about the true internal grain-size
distribution, which should be based on sampling
from deep sediment cores.
2. Reeder et al. (2011) did not provide any sediment
core data in documenting the internal cross-
bedding of the sand dunes. Surface wave geometry
does not routinely imply internal cross-bedding.
Shanmugam 827
Without knowing the type of internal sedimen-
tary structures, the true depositional origin of
sand dunes by traction processes is questionable.
3. Assuming that the surface grain-size distribution
reflects the true internal size distribution, the
origin of dune bedforms composed of fine to
medium-size sands requires current velocities of
50 to 100 cm s
–1
(20–39 in. s
–1
) (Southard,
1975, his figures 2–5). Reeder et al. (2011) did
not provide measured baroclinic current velo-
city data from the study area.
4. Reeder et al. (2011) used the occurrence of sand
dunes on the upper continental slope, away from
the influence of shallow-water tidal forcing, as a
criterion for deep-water internal tidalite origin.
However, shallow-water tidal forcing has been
documented on the deep-water slope environ-
ments (Boyd et al., 2008). Furthermore, Kenyon
et al. (2002) reported sand waves at water depths
of 3200 m (10,499 ft) in the Gulf of Mexico and
attributed their origin to the wind-driven Loop
Current.
5. In the northern South China Sea, where Reeder
et al. (2011) studied sand dunes, migrating sed-
iment waves were previously interpreted to be
deposits of turbidity currents (Damuth, 1979).
Giant sediment waves (5 m [16 ft] in height)
on the continental margin off Nice (southern
France), which are composed of sand and boul-
ders, were ascribed to deposition by sediment
flows (Malinverno et al., 1988).
The exact processes that initiate sediment waves
are still poorly understood (Wynn and Masson,
2008, p. 292). In the midst of these uncertainties,
He et al. (2008, their figure 3) have reinterpreted a
high-resolution seismic profile showing the sym-
metrical wave geometry of muddy sediment waves
from the Rockall Trough area as a product of in-
ternal waves. The problem here is that Howe
(1996, p. 232), who originally studied these sedi-
ment waves (their figure 5), acknowledged the
difficulties associated with interpreting these seismic
wave geometries by concluding, “The mechanisms of
sediment wave formation are not known, although
lee waves propagating over the sinusoidal topography
of the wave forms may have caused the upstream
migration of the waves. The Barra Fan waves may
have been initially constructed by alongslope bot-
tom-current processes with the preexisting sinu-
soidal topography maintained by subsequent domi-
nant downslope turbidity currents.”These sediment
waves were a product of both contour currents and
turbidity currents. He et al. (2008) have injected
unnecessary confusion into this already con-
troversial origin by failing to explain the following
fundamental questions: (1) What depositional mech-
anisms associated with internal waves were respon-
sible for creating wave geometry in muddy sedi-
ments? and (2) What is the seismic criterion that
distinguishes wave geometry created by internal waves
from those created by contour currents or turbidity
currents? At present, no objective criteria exist for
distinguishing wave geometry created by internal
tidal currents from wave geometry created by con-
tour currents or by turbidity currents using seismic
profiles alone (Shanmugam, 2012c).
Systems Tracts
Internal waves were triggered by tropical cyclones
(Nam et al., 2007) and by tsunamis (Santek and
Winguth, 2007). Tropical cyclones also influenced
the generation of internal tides (Davidson and
Holloway, 2003). Empirical data on tropical cy-
clones (meteorological phenomena) and tsunamis
(oceanographic phenomena) from the Indian, At-
lantic, and Pacific Oceans reveal that they are highly
powerful and are common events during the present
sea level highstand (Shanmugam, 2008c). Because
tsunamis are commonly triggered by earthquakes,
no relationship exists between sea level changes and
the timing of tsunamis (Shanmugam, 2008c). For
these reasons, deposits of internal waves and tides
should not be couched into sequence-stratigraphic
systems tracts (lowstand, highstand, and others),
creating unnecessary confusion.
IMPLICATIONS FOR DEEP-MARINE
RESERVOIR SANDS
The only reference to the importance of deposits
of internal waves and internal tides as potential
828 Internal Waves, Internal Tides, and Baroclinic Sands
petroleum reservoirs is by He et al. (2008, p. 42),
who state, “Furthermore, internal-wave and in-
ternal-tide deposits can also form large-scale se-
dimentary bodies, for example, large-scale sed-
iment waves, especially sandy waves. Therefore,
internal-wave and internal-tide deposits should be
favorable exploration targets and potential reser-
voirs in deep-water deposits.”Although no doc-
umented cases of petroleum-producing reservoirs
formed by baroclinic currents exist, their absence
could be attributed to the prevailing lack of cri-
teria to distinguish them from other deposits. At this
embryonic stage of our understanding of baro-
clinic sands, the following provisional comments
on their implications for deep-marine reservoir
sands may be appropriate.
Modern Analogs
In a study of the Kutei Basin in Indonesia, Saller
et al. (2004, p. 42) state, “In the Pleistocene of the
northern Kutei Basin, sand-rich sheetlike strata were
deposited at the toe of slope during the early low-
stand and were fed by conduits that were partially
filled with sand at approximately the same time
(amalgamated channel sands). Three individual
sheet-sand lobes, each approximately 8 by 14 km
(5 by 9 mi) and 50 m (164 ft) thick, coalesce into
the lower basin-floor fan, which is 22 by 22 km (14
by 14 mi) and 60 m (197 ft) thick…. These
sheetlike fan lobes are interpreted to be formed
when turbidites originating on the lowstand delta
front traveled 30 to 40 km (98–131 ft) (sic) ba-
sinward to the basin floor.”These turbidite inter-
pretations with implications for reservoir geometry
and quality are strictly based on seismic geometries
in a sequence-stratigraphic framework. In cases like
this, distinguishing turbidites from internal tidalites
should be based on the pragmatic interpretation of
oceanographic and sedimentologic data from the
Kutei Basin. As noted previously, deep-water
Tertiary sands in the Kutei Basin were interpreted
as turbidites using the Bouma turbidite facies
model by Saller et al. (2006). However, these
turbidite reservoir sands could alternatively be in-
terpreted as tidalites formed by deep-marine tidal
currents using the following empirical data from
the modern Makassar Strait where the Kutei Basin
is located (Shanmugam, 2008a). (1) The Indone-
sian throughflow, which passes through the Ma-
kassar Strait, is stratified along the Makassar sill at a
depth of approximately 680 m (2230 ft) (Gor-
don, 2005). (2) Internal waves (Hatayama, 2004;
Pujiana et al., 2009) and internal tides (Ray et al.,
2005) have been documented in the Makassar
Strait (Figure 1)(Table 2). (3) Satellite imagery
shows well-developed internal waves in the adja-
cent Sulu Sea as well (Figure 6). (4) Nummedal
and Teas (2001) reported the existence of strong
semidiurnal tidal currents flowing along the axis
of a submarine canyon in the Makassar Strait at
a water depth of 993 m (3257 ft), with maxi-
mum current velocities of approximately 40 cm s
–1
(17 in. s
–1
) during the peak of each fortnightly
spring tide. (5) Measurements from two moor-
ings in the Labani Channel recorded velocities in
excess of 50 cm s
–1
(20 in. s
–1
) at 250 m (820 ft)
for baroclinic tidal currents in the Makassar Strait
(Wajsowicz et al., 2003). At these high velocities,
even coarse-grained sand can be reworked by deep-
marine tidal currents. In contrast to these deep-
marine tidal currents, modern turbidity currents
and their velocities have never been documented
from the Kutei Basin in supporting ancient sedi-
ments as turbidites by Saller et al. (2006). The
implication is that the use of modern analogs (i.e.,
uniformitarianism) is imperative in interpreting an-
cient strata in the subsurface.
Stratigraphic Position
Submarine guyots and seamounts of volcanic origin
in the basin-plain environments may complicate the
interpretation of seismic geometries in a sequence-
stratigraphic framework. For example, an Ordovi-
cian seamount in the Roberts Mountains allochthon,
northern Independence Range in Nevada, is se-
lected here for discussion (Watkins and Browne,
1989, their figure 9B). This seamount was the result
of the swelling of the lithosphere and the injection
of basaltic lava onto the ocean floor during the
Middle Ordovician (Figure 17A). Subsequent sub-
sidence and burial of the seamount during the Upper
Ordovician had resulted in a mounded-appearing
Shanmugam 829
feature at the distal position of a turbidite fan
(Figure 17B). If one encounters such a disposition
of deep-marine strata on seismic profiles elsewhere,
the buried seamount could be misinterpreted as a
basin-floor fan based on its stratigraphic position
and mounded geometry (see Vail et al., 1991;
Shanmugam et al., 1995). Such misinterpretations
may result in a surprising reservoir geometry and
quality in petroleum exploration.
Reservoir Geometry
Reservoir geometry is primarily controlled by de-
positional environments. On the basis of deposi-
tional settings, deep-water sands related to internal
waves and tides have been grouped into three
broad categories (He et al., 2008): (1) submarine
channel or canyon (Figure 15A, B, C), (2) un-
channelized continental slope (He et al., 2008, their
figure2D,E,F),and(3)abyssalbasin(Figure 15D).
In a confined submarine canyon environment,
the lateral extent of sand distribution is controlled
by the canyon width. Submarine canyon geom-
etry can be readily recognized on seismic profiles
and on root-mean-square amplitude images. In
the Krishna-Godavari Basin (KG Basin), Bay of
Bengal (India), for example, Pliocene canyons are
sinuous, at least 22 km (14 mi) long, relatively
narrow (500–1000 m [1640–3280 ft] wide), deeply
incised (250 m [820 ft]), and asymmetrically walled
(Shanmugam et al., 2009). Examination of conven-
tional cores shows that sandy debrites and sandy
tidalites occur as sinuous canyon-fill sands. Con-
sidering that internal waves and tides have been
documented in this area (LaFond and Rao, 1954;
Antony et al., 1985), potential exists that the deep-
water sandy tidalites in the submarine canyons of
the KG Basin may be related to internal waves and
tides. Sandy tidalites and related bottom current–
reworked facies exhibit moderate porosity (31–
40%) and permeability (525–6930 md) in the KG
Basin (Shanmugam et al., 2009). Reliance Indus-
tries Limited and Niko Resources discovered gas in
the Pliocene deep-water siliciclastic reservoirs of
the KG Basin in 2002 (Shirley, 2003). These deep-
water sands are major gas-producing reservoirs.
In discussing internal-wave and internal-tide
deposits in abyssal basin environments, He et al.
(2011) proposed an idealized vertical facies model
(Figure 15D). This facies is composed of a basal sandy
turbidite division (layer 1), a middle sandy division
with traction structures formed by reworking by
combined flows and by internal waves and internal
tides (layer 2), and an upper hemipelagic muddy
Figure 17. Conceptual dia-
grams showing two stages of
Ordovician depositional events
in the Roberts Mountains alloch-
thon, northern Independence
Range, Nevada. (A) Swelling of
lithosphere and construction
of seamount (basalt) during
the Middle Ordovician (Early
Llanvirn?). The thick vertical ar-
row points to the second stage.
(B) Subsidence and burial of
seamount during the Upper Or-
dovician (early Caradoc). Litho-
facies descriptions are faithfully
adopted from Watkins and
Browne (1989, their figure 9B,
D), with permission from the
Geological Society of America.
830 Internal Waves, Internal Tides, and Baroclinic Sands
division (layer 3) (Figure 15D). This idealized se-
quence closely mimics the Ta, Tc, and Te divisions
of the Bouma sequence (Bouma, 1962). In describing
these beds, He et al. (2012, p. 368) state, “The beds
are positively graded, and relatively complete clas-
sical Bouma sequences can occasionally be ob-
served….”If so, no difference exists between
turbidites and internal-wave and internal-tide de-
posits. Nevertheless, a fundamental difference is
observed between internal waves and turbidity cur-
rents, which has implications for reservoir geometry
and quality. For example, internal waves and tides
have no capacity for advection of sediment while they
propagate along pycnoclines, bringing new sediment
into an area. That is what turbidity currents do ex-
ceedingly well as sediment-gravity flows (Middleton
and Hampton, 1973). Unlike internal waves, sub-
marine turbidity currents cannot propagate along
the interface between two fluid layers in the ocean.
This is because submarine turbidity currents, as
sediment-gravity flows, have a much higher density
of 1.03 to 2.4 g/cm
3
(Middleton and Hampton,
1973) than sea water (1.03 g/cm
3
). Consequently,
turbidity currents sink to the bottom and flow along
the sea floor. Conventionally, turbidity currents
have been attributed to develop sheetlike sand
geometry in abyssal plain environments (Ricci
Lucchi and Valmori, 1980). In stark contrast,
internal waves and related baroclinic currents are
absent in the abyssal plain environments (Figure 7).
Therefore, internal waves are not important agents
to develop reservoir sands in the abyssal plain settings.
A New Reservoir Facies
In addition to the conventional deep-water envi-
ronments discussed above, towering guyots, sea-
mounts, and mid-ocean ridges are considered to be
sites of internal-wave and internal-tide deposition.
Baroclinic currents tend to rework sands on the up-
per parts of ridges, seamounts, and guyots. As noted
previously, the flat top of the Horizon Guyot is
composed of individual terraces that extend to as
much as 10 km (6 mi) in length and 3 km (1.86 mi)
in width (Lonsdale et al., 1972). These flat surfaces
are suitable environments for reworking of sands by
baroclinic tidal currents (Figure 12). Reworking of
foraminiferal sands on the Horizon Guyot terraces
has developed dune wavelengths of 30 m (98 ft)
and dune heights of 10 m (33 ft) (Figure 13).
Given these impressive dimensions, the develop-
ment of sheetlike geometry of baroclinic sands is
likely. These sheet sands could reach tens of meters
in thickness (Figure 13). However, the challenge of
recognizing these new types of reservoir facies in
the subsurface remains.
If sheet sands with traction structures (e.g.,
cross-bedding, ripple cross-lamination) from an-
cient terraces of guyots were recognized in out-
crops, they are more likely to be the products of
reworking by internal tidal currents than by con-
tour currents or by other bottom currents. This is
because baroclinic currents are the only type that
can propagate along pycnoclines in mid-ocean
depths and can effectively intersect with guyot
tops at any ocean depth. The other bottom cur-
rents, such as contour currents or wind-driven cur-
rents flow along the sea floor. Deep-sea sands with
traction structures formed by current reworking
are potential petroleum reservoirs because of their
high porosity and permeability values caused by
the winnowing away of fines (Mullins et al., 1980).
CONCLUSIONS
1. The significance of deep-marine baroclinic sands,
associated with internal waves and internal tides,
has not received any consideration in the petro-
leum industry. In assisting petroleum geoscien-
tists, empirical data on modern internal waves
and internal tides have been compiled from 51
regions of the oceans of the world.
2. Internal solitary waves (i.e., solitons), the most
common type, are coherent, nonsinusoidal, and
nonlinear in character. They are commonly gen-
erated near the shelf edge (100–200 m [328–
656 ft] in bathymetry) along oceanic pycnoclines.
3. A conceptual oceanographic and sedimento-
logic framework is proposed for explaining de-
position from baroclinic currents on continental
slopes and in submarine canyons. Baroclinic ti-
dal currents, which propagate along oceanic
Shanmugam 831
pycnoclines, can also rework sands on guyot
tops. However, pycnoclines do not intersect the
abyssal (basin) sea floor. Therefore, basin plains
are inapt environments for deposition of baro-
clinic sands.
4. Despite the exhaustive theoretical, experimental,
observational, and numerical analyses of modern
internal waves and tides, no core-based informa-
tion on the origin of primary sedimentary struc-
tures by baroclinic currents in modern marine
environments exists. At present, the criteria for
distinguishing (1) between surface-wave and in-
ternal-wave deposits, (2) between surface-tide
and internal-tide deposits, (3) between inter-
nal-wave and internal-tide deposits, and (4) be-
tween baroclinic sands and turbidite sands are
muddled. Therefore, potential for misinterpret-
ing deep-marine baroclinic sands as something
else is real. Such misinterpretations could result in
surprising economic consequences in the real-
world petroleum exploration.
5. A new kind of reservoir facies, composed of
reworked sands by baroclinic currents, has the
potential to develop sheetlike sand-body geom-
etry with good reservoir quality on the terraces
of towering guyots. Therefore, there is an im-
mediate need to develop criteria to recognize
this long neglected deep-water reservoir facies in
the subsurface.
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