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Kansas Geological Survey, Bulletin 233, p. 63-99


Sedimentary parameters for computer modeling

by
Paul Enos

Department of Geology, University of Kansas

Abstract

The sedimentary parameters that are most important in modeling sedimentary sequences and geometry are accumulation rate, lag time, and accommodation space. Each parameter incorporates several other variables. Accumulation rate is the net result of sediment input and in situ production (for carbonates) less export through bypass or erosion. The appropriate accumulation rate to be chosen from the vast amount of data available will depend on depositional environment, basinal asymmetry, climate, tectonic setting, and the time increment being modeled. Lag time expresses the necessary condition for a transgressive sequence: that the initial sediment accumulation rate is less than the rate of submergence or accommodation. Mechanisms are not well understood; the potential for sediment production in shallow-water carbonate environments, for example, generally exceeds known rates of submergence. Biologic factors may reduce sediment production rates in shallow water, but a physical threshold, such as the wave base, above which accumulation is suppressed, seems more probable. Accommodation space is the increment of room available for sediment accumulation as determined by eustasy, subsidence, and erosion. Subsidence, in turn, incorporates tectonism, isostasy, and physical and chemical compaction. Lag time, compaction rates induced by pressure solution, and the interaction of siliciclastics and carbonates are probably the least constrained variables.

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Introduction

A treatise on improved parameter definitions logically begins with a review of simulation programs to extract the input parameters and critical assumptions, which can then be neatly arrayed in a table. The profusion of available programs, the result of exponential growth from roots in the 1960's [cf. Harbaugh (1966) and Harbaugh and Bonham-Carter (1970)], is such that the table alone would probably exceed the intended length of this article [cf. Aigner et al. (1988), Bice (1988), Bosence and Waltham (1990), Bridge and Leeder (1979), Demicco and Spencer (1989), Harris (1989), Helland-Hanson et al. (1988), Jervey (1988), Koerschner and Read (1989), Lawrence et al. (1990), Lerche et al. (1987), Read et al. (1986), Scaturo et al. (1989), Spencer and Demicco (1989), and Watney et al. (1989)]. The input end of simulation and the other ends as well are reviewed by Kendall et al. (this volume). In this article I focus on sources of precise values for the parameters that most influence the geometry and facies of a simulated sedimentary sequence: accumulation rate, lag time, and accommodation space. Each parameter incorporates other variables. Accumulation rate depends on sediment input and on in situ production in carbonate environments versus sediment removal by erosion and bypassing. Lag time may be the result of biologic or physical thresholds, and accommodation space is the net of eustasy, subsidence, and erosion.

Table 1--Modern sedimentation rates from various depositional settingsa

Area Rate (B)b Period (yr)c Reference
Fluvial environments
*Lower Ohio R., natural levee, 1964 flood 460,000 1 Bridge & Leeder, 1979 (Alexander & Prior, 1971)
*Ohio R. floodplain, 1937 flood average 70,000 1 Bridge & Leeder, 1979 (Mansfield, 1938)
*Ohio R. floodplain, 1937 flood range 3,000-560,000 1 Bridge & Leeder, 1979 (Mansfield, 1938)
Lower Ohio R., natural levee 16,000 40 Bridge & Leeder, 1979 (Alexander & Prior, 1971)
Lower Ohio R., natural levee 10,000 750 Bridge & Leeder, 1979
Lower Ohio R., accretionary ridge 6,000 1,000 Bridge & Leeder, 1979
Ohio R. floodplain 4,500 150 Schindel, 1980 (Moore, 1971)
*Lower Ohio R., accretionary ridge, 1964 flood 3,200 1 Bridge & Leeder, 1979 (Alexander & Prior, 1971)
Lower Ohio R., swale 1,900 1,000 Bridge & Leeder, 1979
Lower Ohio R., accretionary ridge 270 1,000 Bridge & Leeder, 1979
Yuba R., California 100,000 1 Kukal, 1971
Sacramento R., California 75,000 1 Kukal, 1971
Cimarron R., Maryland, floodplain 51,000 12 Schindel, 1980 (Schumm & Lichty, 1963)
Western Run, Maryland, floodplain 16,300 50 Schindel, 1980 (Costa, 1973)
Nile R., floodplain 9,000 1 Kukal, 1971
Nile R., floodplain, range 9,100-12,200 1,000 Bridge & Leeder, 1979 (Leopold et al., 1964)
Delaware R. floodplain 140-1,150 6,000 Schindel, 1980 (Ritter et al., 1973)
Indus R. 200 4,500 Kukal, 1971
Wisconsin valley floodplain 1,000 6,070 Bridge & Leeder, 1979 (Knox, 1972)
Wisconsin valley floodplain 350 6,040 Bridge & Leeder, 1979 (Knox, 1972)
Blockhouse Creek, Wisconsin, floodplain 150-380 6,000 Bridge & Leeder, 1979 (Knox, 1972)
Little Tallahatchie R., Mississippi, natural levee 47,000-65,000 8 Bridge & Leeder, 1979 (Ritchie et al., 1975)
Little Tallahatchie R., natural levee 13,000-20,000 31 Bridge & Leeder, 1979 (Ritchie et al., 1975)
Little Tallahatchie R., crevasse splay 28,000-34,000 8 Bridge & Leeder, 1979 (Ritchie et al., 1975)
Little Tallahatchie R., crevasse splay 27,000 31 Bridge & Leeder, 1979 (Ritchie et al., 1975)
Little Tallahatchie R., abandoned channels 9,000-28,000 8 Bridge & Leeder, 1979 (Ritchie et al., 1975)
Little Tallahatchie R., abandoned channels 7,000-10,000 31 Bridge & Leeder, 1979 (Ritchie et al., 1975)
Beatton R., British Columbia, floodplain 1,000-61,000 500 Bridge & Leeder, 1979 (Nanson, 1977)
Bobr & Strzegomka R., USSR, floodplains 1,000-5,000 ≈10,000 Bridge & Leeder, 1979 (Teisseyre, 1977)
South Carolina Piedmont rivers, floodplain 8,000 150 Bridge & Leeder, 1979 (Happ, 1945)
Buck Run, floodplain 650 1,450 Bridge & Leeder, 1979 (Leopold et al., 1964)
Tigris & Euphrates, floodplain 200 5,000 Bridge & Leeder, 1979 (Leopold et al., 1964)
Cheyenne R., Wyoming, floodplain 41,000-61,000 60 Bridge & Leeder, 1979 (Leopold et al., 1964)
Dry Creek, Nebraska, floodplain 8,600 500 Bridge & Leeder, 1979 (Brice, 1966)
Upper Dry Creek, Nebraska, floodplain 4,600-5,500 33 Bridge & Leeder, 1979 (Brice, 1966)
Well Canyon, Nebraska, floodplain 15,500-20,000 40 Bridge & Leeder, 1979 (Brice, 1966)
Medicine Creek, Nebraska, floodplain 83,000 22 Bridge & Leeder, 1979 (Brice, 1966)
Medicine Creek, drainage basin average 25,000 22-500 Bridge & Leeder, 1979 (Brice, 1966)
Chemung R., New York, floodplain 4,600   Bridge & Leeder, 1979 (Nelson, 1966)
*Bijou Creek, Colorado, overbank, 1965 flood 61,000-3,600,000 1 Schindel, 1980 (McKee et al., 1967)
*Missouri R., levees, 1881 flood 1.22-1.83 x 106 1 Bridge & Leeder, 1979 (Leopold et al., 1958)
*Kansas R., floodplain, 1951 flood 29,000 1 Bridge & Leeder, 1979 (Leopold et al., 1958)
*Farmington R., Connecticut, floodplain, 1955 flood 15,000 1 Bridge & Leeder, 1979 (Wolman & Eiler, 1958)
*Connecticut R., floodplain, 1936 flood 35,000 1 Bridge & Leeder, 1979 (Jahns, 1947)
*Connecticut R., floodplain, 1938 flood 22,000 1 Bridge & Leeder, 1979 (Jahns, 1947)
*Connecticut R., banks, 1936 flood 259,000 1 Bridge & Leeder, 1979 (Jahns, 1947)
*Connecticut R., banks, 1938 flood 173,000 1 Bridge & Leeder, 1979 (Jahns, 1947)
*Connecticut R., tributary banks, 1936 flood 200,000 1 Bridge & Leeder, 1979 (Jahns, 1947)
*Connecticut R., tributary banks, 1938 flood 107,000 1 Bridge & Leeder, 1979 (Jahns, 1947)
*Connecticut R., artificial levees, 1936 flood 91,000 1 Bridge & Leeder, 1979 (Jahns, 1947)
*Connecticut R., artificial levees, 1938 flood 43,000 1 Bridge & Leeder, 1979 (Jahns, 1947)
*Ob' R., USSR, point bar, 1969 flood ≤1,500,000 1 Bridge & Leeder, 1979 (Velikanov & Yamykh, 1970)
*Ob' R., crevasse splay & levee, 1969 flood ≤600,000 1 Bridge & Leeder, 1979
*Ob' R., USSR, flood basin, 1969 flood 200-30,000 1 Bridge & Leeder, 1979
San Joaquin River, California, (Holocene) 15,000 ≈10,000 Bull, 1972
Mississippi R., floodplain 1,400 30,000 Bridge & Leeder, 1979 (Fisk, 1944)
Upper Mississippi R., artificial backwater 25,000-35,000 20 Bridge & Leeder, 1979 (McHenry et al., 1976)
*Mississippi R., point bar, 1973 flood 860,000 1 Bridge & Leeder, 1979 (Kesel et al., 1974)
*Mississippi R., point bar, range 130,000-3,000,000 1 Bridge & Leeder, 1979 (Kesel et al., 1974)
*Mississippi R., natural levee, 1973 flood 530,000 1 Bridge & Leeder, 1979 (Kesel et al., 1974)
*Mississippi R., natural levee, 1973 flood 100,000-840,000 1 Bridge & Leeder, 1979 (Kesel et al., 1974)
*Mississippi R., levee back, 1973 flood 125,000 1 Bridge & Leeder, 1979 (Kesel et al., 1974)
*Mississippi R., levee back, 1973 flood 60,000-270,000 1 Bridge & Leeder, 1979 (Kesel et al., 1974)
*Mississippi R., abandoned channels, 1973 flood 60,000 1 Bridge & Leeder, 1979 (Kesel et al., 1974)
*Mississippi R., abandoned channels, 1973 flood,range 40,000-90,000 1 Bridge & Leeder, 1979 (Kesel et al., 1974)
*Mississippi R., backswamp, 1973 flood 11,000 1 Bridge & Leeder, 1979 (Kesel et al., 1974)
*Mississippi R., backswamp, 1973 flood 5,000-25,000 1 Bridge & Leeder, 1979 (Kesel et al., 1974)
Mississippi River (Miocene) 32-53 20 x 106 Rainwater, 1966
Diablo Range, California, Holocene alluvial fan 26,000 ≈10,000 Bull, 1972
Eolian environments
Southern Peru 2,000,000 3 Bigarella, 1972
Sahara Desert, Algeria 6,500 4,000 Galloway & Hobday, 1983 (Wilson, 1973)
Southern Sahara Desert 800-1,700 12,000 Breed et al., 1979
Grand Erg Oriental, Algeria, average 19 1,350,000 Breed et al., 1979
Grand Erg Oriental, Algeria, max. 87 1,350,000 Breed et al., 1979
Kalahari Desert 300-3,300 ≈10,000 Breed et al., 1979
Cerchen Desert, PRC, average 60,000 1,500 Breed et al., 1979
Navajo Sandstone (E. Jurassic), USA 53 17 x 106 Galloway & Hobday, 1983
Loess 200-1,000   Kukal, 1971
Loess, central Alaska 15-193   Kukal, 1971 (Péwé, 1968)
Lacustrine environments
Lacustrine, average 3,000   Kukal, 1971
Vierwaldstättersee, Switzerland, calcareous clays 10,400-31,700   Kukal, 1971
Vierwaldstättersee, Switzerland 3,500-5,000   Schindel, 1980 (Schwarzacher, 1975)
Brienz, Switzerland, calcareous clays 31,700   Kukal, 1971
Léman, Switzerland 1,200 ≈150 Schindel, 1980 (Krishnaswami et al., 1971)
Lunz, [Léman?], mouth of Rhone 17,900   Kukal, 1971
Lunz, [Léman?], average 2,500   Kukal, 1971
Lunz, Austria, average 1,800   Schindel, 1980 (Schwarzacher, 1975)
Salton Sea, California, saline 5,000-14,000 ≈50 Schindel, 1980 (Amal, 1961)
Wallensee, Switzerland, calcareous clays 11,300   Kukal, 1971
Onega, USSR, clays 7,100   Kukal, 1971
Ladoga, USSR, clays 6,120   Kukal, 1971
Trout and Mendota, Wisconsin 6,000 <100 Schindel, 1980 (Bruland et al., 1975)
Trout Lake, Wisconsin 4,000 ≈100 Schindel, 1980 (Koide et al., 1972)
Olof Jone Damm, Sweden, peat 5,300   Kukal, 1971
Shinji, Japan 3,000-5,000 <100 Schindel, 1980 (Matsumoto, 1975)
Shinji, Japan 1,200 9,500 Schindel, 1980 (Mizuno et al., 1972)
North German lakes, marl 1,000-3,000   Kukal, 1971
Swedish lakes, gyttja 1,000-2,000   Kukal, 1971
Pavin, France 1,300 ≈100 Schindel, 1980 (Krishnaswami et al., 1971)
Tahoe, USA 1,000 ≈100 Schindel, 1980 (Koide et al., 1972; Bruland et al., 1975)
Titicaca, Bolivia 1,000   Schindel, 1980 (Bruland et al., 1975)
Zürich, Switzerland 700   Kukal, 1971
Neuchâtel, Switzerland, varved calcareous clays 700   Kukal, 1971
Maxinkuckee, Canada, marl, eutrophic 300   Kukal, 1971
Great Lakes, N. America, varved mud 150   Kukal, 1971
Michigan, varved calcareous clays 3,000   Kukal, 1971
Michigan 100-1,200 <100 Schindel, 1980 (Robbins & Edgington, 1975)
Superior 100-600 <100 Schindel, 1980, (Bruland et al., 1975)
Constance (Bodensee), mouth of Rhine 22,400 15,000 Müller & Gees, 1970
Constance (Bodensee) 1,500-6,000 15,000 Müller & Gees, 1970
Diatomite (average) 300-1,000   Kukal, 1971
Varved glacial lakes
Weistriztal, Czechoslovakia 60,000-100,000   Reineck & Singh, 1975 (Schwarzbach, 1940)
Burks Falls, Ontario 2,000-17,000 620 Antevs, 1925
Espanola, Ontario 1,000-83,000 985 Antevs, 1925
Tishaming, Ontario 4,000-65,000 1,335 Antevs, 1925
Huntsville, Ontario 2,000-45,000 760 Antevs, 1925
Bracebridge, Ontario 3,000-92,000 511 Antevs, 1925
Bracebridge, Ontario, average 10,900 112 Antevs, 1925
Ancient lakes
Lake Bonneville, Pleistocene, Utah 1,800 106 Feth, 1964; Picard & High, 1972
Unita Formation, Eocene, Wyoming 5.7 13.3 x 106 Feth, 1964; Picard & High, 1972
Green River Formation, Eocene, Wyoming and Colorado 150 4 x 106 Feth, 1964; Picard & High, 1972
Flagstaff Limestone, Paleocene-Eocene, Utah 22-110 2.75 x 106 Feth, 1964; Picard & High, 1972
Todilto Limestone, L. Jurassic, New Mexico 380 20,000 Feth, 1964; Picard & High, 1972
Lockatong Formation, L. Triassic, New Jersey 225 5.1 x 106 Van Houten, 1964; Picard & High, 1972
Deltaic environments
Delta topsets, average 15,000-20,000   Kukal, 1971
Mississippi 2,740,000 4 d Kukal, 1971
Mississippi, channel-mouth bar 500,000 100 Schindel, 1980 (Coleman, 1976)
Mississippi, channel-mouth bar 340,000 195 Schindel, 1980 (Gould, 1970)
Mississippi, delta front 300,000-450,000 1 Kukal, 1971
Mississippi, prodelta 60,000-300,000 1 Kukal, 1971
Mississippi, subaerial average 170,000-200,000 600 Schindel, 1980 (Coleman, 1976)
Mississippi, offshore 200,000 100 Schindel, 1980 (Coleman, 1976)
Mississippi, crevasse splay 30,000-100,000 150 Schindel, 1980 (Coleman, 1976)
Mississippi, adjacent shelf 45,000 1 Kukal, 1971
Mississippi, premodern lobes, average 20,000-25,000 1,000 Schindel, 1980 (Coleman, 1976)
Mississippi, interdistributary bay 19,600 120 Elliott, 1978 (Gagliano & Van Beck, 1970)
Mississippi, Sale Sypremort lobe 8,300-12,500 1,200 Schindel, 1980 (Coleman, 1976)
Mississippi, "maximum" 10,000 11,000 Lisitzin, 1972
Mississippi, submarine, nearshore 8,200 <100 Schindel, 1980 (Shokes & Presley, 1976)
Mississippi, submarine, cont. shelf 6,100 <100 Schindel, 1980 (Shokes & Presley, 1976)
Mississippi, submarine, cont. slope 400 <100 Schindel, 1980 (Shokes & Presley, 1976)
Mississippi River (Miocene) 325 20 x 106 Rainwater, 1966
Mississippi River (Miocene), prodelta 220 20 x 106 Rainwater, 1966
Colorado, Texas 4,064,000 6 Kanes, 1970
Po, Italy, at shoreline, average 465,000 25 Nelson, 1970
Po, at shoreline, range 268,000-653,000 19-45 Nelson, 1970
Rhône 400,000 1 Kukal, 1971
Rhône 700   Schindel, 1980 (Schwarzacher, 1975)
Rhône, river mouth, 50 m depth 350,000 1 Oomkens, 1970
Rhône, mouth of Grand Rhône 14,000 ≈5,000 Oomkens, 1970
Rhône, mouth of Petit Rhône 7,600 ≈5,000 Oomkens, 1970
Rhône, shoreline 2,000 11,000 Lisitzin, 1972
Rhône, 45 km offshore 6,000 11,000 Lisitzin, 1972
Rhône, 75 km offshore 1,000 11,000 Lisitzin, 1972
Rhine delta, Lake Constance (Bodensee) 2,500,000 10 Müller, 1966
Rhine delta, Lake Constance, average 262,800 50 Müller, 1966
Nile, subaerial portion 10,000 1 Kukal, 1971
Nile 660   Schindel, 1980 (Schwarzacher, 1975)
Fraser, Canada 50,000-300,000   Kukal, 1971
Volga, Caspian Sea 5,000-70,000   Kukal, 1971
Tana River, Japan 30,000-70,000 10 Schindel, 1980 (Ambe, 1972)
Alamo River, Salton Sea, USA 50,000 33 Schindel, 1980 (Amal, 1961)
Amu Darya River, Aral Sea, USSR 25,000   Kukal, 1971
Orinoco, Venezuela 5,000-6,000 11,000 Lisitzin, 1972
Sabine, Texas 2,930 5,200 Nelson & Bray, 1970
Guadalupe, Texas 2,100 2,000 Donaldson et al., 1970
Rud Hilla, Persian Gulf 800-5,000 <6,000 Schindel, 1980 (Melguen, 1973)
Columbia River, Washington shelf 1,300-3,900 ≈100 Schindel, 1980 (Nittrouer et al., 1979)
Huang-He, PRC 1,500   Kukal, 1971
Don, Sea of Azov, USSR, subaerial portion 1,220 1 Kukal, 1971
Malaysia, tide-dominated delta 1,000 100 Galloway & Hobday, 1983 (Coleman et al., 1970)
Bengal cone (Ganges prodelta) 62 10.2 x 106 Moore et al., 1974
Tidal flats, coastal wetlands, and beaches
Tidal flats
Jade Busen, Germany 1,450,000 8 d Kukal, 1971 (Reineck, 1960)
Jade Busen, Germany 11,500 4 Kukal, 1971 (Reineck, 1960)
Jade Busen, Germany 2,200 1,900 Kukal, 1971 (Reineck, 1960)
The Wash, UK 16,000-80,000 9 mo Schindel, 1980 (Evans, 1965)
Netherlands 10,000-20,000 1 Kukal, 1971
Laguna Madre, Texas 250-5,000 2,500 Schindel, 1980 (Miller, 1975)
Boundary Bay, British Columbia 5,000 20 Schindel, 1980 (Kellerhals & Murray, 1969)
Boundary Bay, British Columbia 420 4,500 Schindel, 1980 (Kellerhals & Murray, 1969)
Beaches
Chenier beaches, SW Louisiana 6,300-21,200 400-2,200 Reineck & Singh, 1975 (Gould and McFarlan, 1959)
Padre Island, Texas, barrier beach 2,100 4,000 Reineck & Singh, 1975 (Fisk, 1959)
Galveston Island, Texas, barrier beach 2,860 3,500 Bernard et al., 1962
Fire Island Inlet, New York 103,000 115 Kumar & Sanders, 1974
Wangerooge Inlet, North Sea 88,000 68 Reineck & Singh, 1975 (Reineck, 1958)
Nayarit, Mexico 44,000 205 Reineck & Singh, 1975 (Curray et al., 1969)
Christchurch Formation (Holocene), New Zealand, offshore sand 2,700-3,500 5,535 Suggate, 1968
Wetlands, salt marshes
Long Island Sound, USA 4,700-6,300 <100 Schindel, 1980 (Annentano & Woodwell, 1975)
Denmark 3,600 30 Schindel, 1980 (Schou, 1967)
Connecticut 2,000-6,500 6 Schindel, 1980 (Harrison & Bloom, 1974)
Farm River, Connecticut 1,600 200 Schindel, 1980 (McCaffrey, 1977)
Klang River (Malaysia) 1,000   Schindel, 1980 (Coleman, 1976)
SW Louisiana, salt marsh & lagoon 5,500-27,500 400-1,800 Reinick & Singh, 1975 (Gould and McFarlan, 1959)
Wetlands, peat deposits
UK, Littoral 9,800   Kukal, 1971
Schwabia, S. Germany, high moors 1,500-1,800 1 Kukal, 1971
North American 550   Kukal, 1971
Bomeo (Kalimantan), coastal swamps 4,250 4,000 Galloway & Hobday, 1983 (Stach et al., 1975)
Everglades, Holocene coastal swamp 1,200 3,460 Spackman et al., 1964
Olof Jone Damm, Sweden, fresh water 5,300   Kukal, 1971
Bays, lagoons, and estuaries
Texas, lagoon 14,300 290 Schindel, 1980 (Moore, 1955)
Texas, lagoon 9,100 68 Schindel, 1980 (Shepard, 1953)
Texas, lagoon, clay and eolian sand 3,800   Kukal, 1971
San Antonio Bay, Texas 3,750 100 Donaldson et al., 1970 (Shepard & Moore, 1960)
Texas, lagoon 2,300 9,300 Schindel, 1980 (Shepard & Moore, 1955)
Padre Island, Texas, lagoon 1,900 4,000 Reineck & Singh, 1975 (Fisk, 1959)
Great Bay, USA, estuary 1,600-7,800 ≈100 Schindel, 1980 (Capuzzo & Anderson, 1973)
Long Island Sound, USA 6,000 30 Schindel, 1980 (Thomson & Turekian, 1973)
Long Island Sound 1,000-7,000 <100 Schindel, 1980 (Benninger et al., 1977)
Long Island Sound 500-1,000 10,000 Schindel, 1980 (Benninger et al., 1977)
Mobile Bay, Alabama 5,600 115 Schindel, 1980 (Ryan & Goodell, 1972)
Mobile Bay, Alabama 1,640 6,000 Schindel, 1980 (Ryan & Goodell, 1972)
Firth of Clyde, Scotland 5,000 12,000 Schindel, 1980 (Kuenen, 1950)
Firth of Clyde, Scotland, clay 2,400-3,000   Kukal, 1971
James River, Virginia, estuary 1,500-3,000 75 Schindel, 1980 (Nichols, 1972)
Sea of Azov, USSR, estuary 900-2,400 11,000 Lisitzin, 1972
Hampton, New Hampshire, estuary 1,000-2,300 11,000 Schindel, 1980 (Keene, 1970)
Kiel Bay, Germany, sand & silt 1,500-2,000   Kukal, 1971
Drammens Fjord, Sweden 1,500 12,000 Schindel, 1980 (Kuenen, 1950)
San Francisco Bay, USA 300-1,300 ≈2,500 Schindel, 1980 (Story et al., 1966)
Gulf of California, Mexico 1,000 12,000 Schindel, 1980 (Kuenen, 1950)
Gulf of California, clay, diatomite 600-1,000   Kukal, 1971
Gulf of Paria, Venezuela, clay 0-10,000 700 Kukal, 1971 (van Andel & Postma, 1954)
Kara-Bougas-Gol, Caspian Sea, clay & salt 500-700   Kukal, 1971
Inland seas
Black Sea (Messinian), carbonates, some pebbly mudstone 1,030 800,000 Hsü, 1978
Black Sea, terrigenous clastics 1,000 15,000 Schindel, 1980 (Stoffers et al., 1978)
Black Sea, dolomitic varves 900 81 Schindel, 1980 (Ross et al., 1978)
Black Sea, terrigenous & diatom mud 600 125,000 Hsü, 1978
Black Sea, terrigenous & diatom mud 500 500,000 Hsü, 1978
Black Sea, lacustrine carbonates 310 1.1 x 106 Hsü, 1978
Black Sea, terrigenous mud 100-400 11,000 Lisitzin, 1972
Black Sea 50-400 7,000 Schindel, 1980 (Ross et al., 1970)
Black Sea, coccolith ooze 100-300 3,000 Schindel, 1980 (Stoffers et al., 1978)
Black Sea 200   Kukal, 1971
Black Sea 200 12,000 Schindel, 1980 (Kuenen, 1950)
Black Sea, brackish sapropel 100 4,000 Schindel, 1980 (Stoffers et al., 1978)
Black Sea, marine, terrigenous mud 100 11,000 Hsü, 1978
Black Sea (Pliocene) lacustrine chalk 54 3.45 x 106 Hsü, 1978
Black Sea (pelagic) 10-40 11,000 Lisitzin, 1972
Black Sea (Miocene) black shale 26 4 x 106 Hsü, 1978
Mediterranean, Baleric abyssal plain; hemipelagic & turbidites 160-520 20,000 Rupke, 1975
Mediterranean, Tyrrhenian Sea, calcareous and diatom clay 100-500 12,000 Schindel, 1980 (Kuenen, 1950)
Mediterranean, terrigenous turbidites 300 10,000 Cita et al., 1978
Mediterranean, ooze & eolian 200   Kukal, 1971
Mediterranean, deep basin 150 1.9 x 106 Cita et al., 1978
Mediterranean, calcareous ooze 100   Kukal, 1971
Mediterranean, average 25-90 3 x 106 Cita et al., 1978
Mediterranean, pelagic oozes 50 ≈10,000 Cita et al., 1978
Mediterranean, ridges & basin margins 25-50 1.9 x 106 Cita et al., 1978
Mediterranean, Adriatic Sea, shells (lag) 10   Kukal, 1971
Mediterranean, pelagic (condensed) 1 1.9 x 106 Cita et al., 1978
Caspian Sea, mouth of Kura River 6,000 7,000 Lisitzin, 1972
Caspian Sea, pelagic 200-600 7,000 Lisitzin, 1972
Caspian Sea, calcareous clays 100-180   Kukal, 1971
Baltic Sea 200-2,000 10,000 Schindel, 1980 (Alhonen, 1966)
Baltic Sea, black organic clays 300   Kukal, 1971
Persian Gulf (eastern basin), terrigenous 410 9,000 Schindel, 1980 (Seibold et al., 1973)
Persian Gulf (central basin), terrigenous and carbonate 70 9,000 Schindel, 1980 (Seibold et al., 1973)
Sea of Okhotsk, W. Pacific, shelf depression and base of slope 90-250 11,000 Lisitzin, 1972
Sea of Okhotsk, central shelf 9-45 11,000 Lisitzin, 1972
Gulf of Mexico, upper slope, sandy, silty clays 70   Kukal, 1971
Gulf of Mexico, lower slope, silty clays 50   Kukal, 1971
Gulf of Mexico, basin floor, calcareous clays 40   Kukal, 1971
Milford Sound (New Zealand), sandy silts 12.5   Kukal, 1971
Terrigenous shelf deposits
North American shelf 0-400   Kukal, 1971
New Jersey shelf, USA, sand 950-1,300 8,000 Swift et al., 1984
Barents Sea, Arctic Ocean, clays 8-40   Kukal, 1971
North Sea, gale 'Adolph-Bennpohl', sand 4,200,000 36 h Gadow & Reineck, 1969
North Sea, storm, March 1967, sand 1,100,000 84 h Gadow & Reineck, 1969
Nompho, Korea, mud 1,500,000 4 Lisitzin, 1972
Antarctic shelf, sand 20-60 10,000 Lisitzin, 1972
Antarctic shelf, mud 200-300 10,000 Lisitzin, 1972
Indochinese shelf 50-200   Kukal, 1971
North Sea sand waves 2,200-4,300 10,000 Houbolt, 1968
High Island, Gulf of Mexico, nearshore sands 1,050-1,760 5,200 Nelson & Bray, 1970
High Island, Gulf of Mexico, nearshore muds 470 5,200 Nelson & Bray, 1970
Shallow-water carbonates
Individual coral growth rates
Range, coral growth 850-150,000   Schlager, 1981
Massive coral 4,000   Kukal, 1971
Pristatophyllum, Devonian 2,000-6,200 33 Faul, 1943
Atlantic corals
Corals, Florida Bay, leeward 570 528 Kukal, 1971
Montastrea annularis
inshore, patch reef, <6 m depth, Florida 8,200 50 Shinn, Lidz et al., 1989 (Hudson, 1981)
offshore, >6 m depth, Florida 6,300 50 Shinn, Lidz et al., 1989 (Hudson, 1981)
platform margin, <3 m depth, Florida, windward 11,200 50 Shinn, Lidz et al., 1989 (Hudson, 1981)
Key West, Florida 2,800-5,800 15 Weber & White, 1977
Florida, 0-5 m 6,000 3 Huston, 1985 (Vaughan, 1915)
Virgin Islands, 0 m 9,170 ± 1,330   Baker & Weber, 1975
Virgin Islands, 5 m 9,950 ± 1,430   Baker & Weber, 1975
Virgin Islands, 9 m 10,410 ± 1,240   Baker & Weber, 1975
Virgin Islands, 13.5 m 9,690 ± 1,360   Baker & Weber, 1975
Virgin Islands, 18 m 6,540 ± 3,560   Baker & Weber, 1975
Virgin Islands, 22.5 m 2,060 ± 540   Baker & Weber, 1975
Virgin Islands, 27 m 1,560 ± 200   Baker & Weber, 1975
Virgin Islands, 2 m, leeward 7,600 ± 330 4 m Gladfelter et al., 1978
Virgin Islands, forereef, windward 10 m, 7,600 ± 820 4 m Gladfelter et al., 1978
Jamaica, 0-1 m 6,950 3 Huston, 1985
Jamaica, 5 m 7,400 ± 3,100 3 Huston, 1985
Jamaica, 10 m 7,430 ± 1,920 3 Huston, 1985
Jamaica, 10 m 6,680 ± 2,000 2 Dustan, 1979
Jamaica, 20 m 1,770 ± 700 3 Huston, 1985
Jamaica, 28 m 1,700 ± 400 2 Dustan, 1979
Jamaica, 30 m 1,750 ± 330 3 Huston, 1985
Jamaica, 45 m 1,630 ± 1,200 2 Dustan, 1979
Curacao, 6-15 m 6,300-7,800   Huston, 1985 (Bak, 1976)
Belize 12,000   Weber & White, 1977
Caribbean 3,000-12,000   Weber & White, 1977
Caribbean 6,000   Ghiold & Enos, 1982 (Macintyre & Smith, 1974)
Pleistocene, Florida 5,000 400 Shinn, Lidz et al., 1989
Pleistocene, Florida 5,200 31 Landon,1975
Montastrea cavernosa
Jamaica, 10 m 3,600 ± 1,900 3 Huston, 1985
Jamaica, 20 m 6,840 ± 2,670 3 Huston, 1985
Jamaica, 30 m 4,100 ± 1,390 3 Huston, 1985
Key West 2,250-4,050 22 Weber & White, 1977
Florida 3,200-5,700 3 Ghiold & Enos, 1982 (Vaughn, 1915)
Acropora palmata
Caribbean 100,000-120,000 1 Shinn, Lidz et al., 1989 (Lewis et al., 1968)
Florida, <5 m 25,000-40,000 3 Huston, 1985 (Vaughan, 1915)
Curacao, <5 m 88,000   Huston, 1985 (Bak, 1976)
Virgin Islands, 1/2 m, leeward 56,900 ± 4,100 2 m Gladfelter et al., 1978
Virgin Islands, 1/2 m, windward 68,500 ± 6,900 2 m Gladfelter et al., 1978
Virgin Islands, 9 m , windward 77,000 ± 6,900 2 m Gladfelter et al., 1978
Acropora cervicornis
Jamaica, windward 266,000 ± 129,000 1 Buddemeier & Kinzie, 1976 (Lewis et al., 1968)
Jamaica, <5 m, windward 109,000-159,000 1 Tunnicliffe, 1983
Jamaica, 6-15 m, windward 80,000-140,000 1 Tunnicliffe, 1983
Jamaica, 25 m, windward 92,000-148,000 1 Tunnicliffe, 1983
Virgin Islands, 10 m, windward 71,000 ± 6,500 2 m Gladfelter et al., 1978
Florida, <5 m, windward 105,100 ± 16,500 1 Shinn, 1966
Florida, 1 m, leeward 45,700 ± 18,400 9 m Shinn, 1966
Florida, <5 m 40,000-45,000 3 Huston, 1985 (Vaughan, 1915)
Barbados, leeward 145,000 ± 559,000 1 Buddemeier & Kinzie, 1976 (Lewis et al., 1968)
Colpophyllia natans
Jamaica, 5 m 9,000 ± 1,100 3 Huston, 1985
Jamaica, 10 m 8,100 ± 1,400 3 Huston, 1985
Jamaica, 20 m 4,200 ± 850 3 Huston, 1985
Diploria spp. 3,500-10,000 3 Ghiold & Enos, 1982 (Vaughn, 1915)
Diploria stringosa, Bermuda 3,300-3,500 50 Dodge & Vaisnys, 1980
D. labyrinthiformis, Florida 3,500 ± 600 3-27 Ghiold & Enos, 1982
Solenastrea bournoni, Florida Bay, leeward 8,900 100 Shinn, Lidz et al., 1989
Porites porites
Florida 14,000-17,000 6 Landon,1975
Caribbean 21,000-36,000   Landon, 1975 (Lewis et al., 1968)
Pleistocene, Florida 10,500 6 Landon, 1975
Poritesfurcata, Florida & Bahamas 9,000-23,000 3 Ghiold & Enos, 1982 (Vaughn, 1915)
Porites astreoides
Jamaica, 0-1 m 5,030 ± 560 3 Huston, 1985
Jamaica, 5 m 5,000 ± 1,500 3 Huston, 1985
Jamaica, 10 m 3,300 ± 770 3 Huston, 1985
Jamaica, 20 m 2,700 ± 220 3 Huston, 1985
Jamaica, 30 m 2,300 ± 250 3 Huston, 1985
Virgin Islands, 2 m 3,450 ± 320 8 m Gladfelter et al., 1978
Virgin Islands, 10 m 3,000 ± 120 3 m Gladfelter et al., 1978
Florida 4,300   Ghiold & Enos, 1982 (Kissling, 1977)
Florida, Bahamas 5,700-13,000 3 Ghiold & Enos, 1982 (Vaughn, 1915)
Dendrogyra cylindrus, Florida 5,000 1 Shinn, Lidz et al., 1989
Siderastrea sidera
Jamaica, 10 m 7,150 3 Huston, 1985
Jamaica, 20 m 3,000 ± 800 3 Huston, 1985
Jamaica, 30 m 3,100 3 Huston, 1985
Florida 2,200-2,700 22 Landon,1975
Pleistocene, Florida 1,500 19 Landon, 1975
Faviafragum, Florida 2,900-3,800 3 Ghiold & Enos, 1982 (Vaughn, 1915)
Manicina sp. Florida 2,500-8,700 3 Ghiold & Enos, 1982 (Vaughn, 1915)
Agaricia argaricites, Jamaica, 0-30 m 1,110 ± 270 ≥3 Huston, 1985
Agaricia sp, Florida 3,800 3 Ghiold & Enos, 1982 (Vaughn, 1915)
Pacific corals
Acropora spp 85,000-226,000   Buddemeier & Kinzie, 1976
Acropora spp, Samoa, 0-5 m 4,000-185,000   Huston, 1985 (Mayor, 1924)
A. abrantoides, Yapp 125,000-130,000   Huston, 1985 (Tamura & Hada, 1932)
A. pulchra 101,000-172,000   Huston, 1985 (Tamura & Hada, 1932)
Astreapora myriophythalma
Enewetak, 6-15 m 7,500-13,000   Huston, 1985 (Buddemeier et al., 1974)
Enewetak, 16-25 m 5,000-5,500   Huston, 1985 (Buddemeier et al., 1974)
Pocillopora spp., Samoa, 0-5 m 7,000-35,000   Huston, 1985 (Mayor, 1924))
P. damicornis 13,900-27,800   Buddemeier & Kinzie, 1976
Guam, 0-5 m 33,300   Huston, 1985 (Neudecker, 1977)
Guam, 6-15 m 36,700   Huston, 1985 (Neudecker, 1977)
Guam, >25 m 18,100   Huston, 1985 (Neudecker, 1977)
Panama, 3 m 39,600 ± 1,500   Glynn, 1976
Panama, 6 m 33,600 ± 2,100   Glynn, 1976
Panama, 0-15 m 44,300-59,300   Huston, 1985 (Wellington, 1982)
P. eydouxi, Enewetak, 0-5 m 50,000   Huston, 1985 (Buddemeier et al., 1974)
Psammocora sp., Enewetak, 0-5 m 30,000   Huston, 1985 (Buddemeier et al., 1974)
Pavona sp., Samoa, 0-5 m 32,000   Huston, 1985 (Mayor, 1924)
P. clavus, Panama, 0-5 m 15,500-23,000   Huston, 1985 (Wellington, 1982)
Panama, 6-15 m 12,000-19,000   Huston, 1985 (Wellington, 1982)
P. gigantea, Panama, 0-5 m 10,000-19,500   Huston, 1985 (Wellington, 1982)
Panama 6-15 m 8,000-17,000   Huston, 1985 (Wellington, 1982)
Fungiafungites, Enewetak, 6-15 m 10,000-12,000   Huston, 1985 (Buddemeier et al., 1974)
Porites spp, 0-5 m 7,000-48,500   Huston, 1985 (various)
P. lutea
Enewetak, 0-5 m 5,000-13,500   Huston, 1985 (Buddemeier et al., 1974)
Enewetak, 6-15 m 5,000-11,000   Huston, 1985 (Buddemeier et al., 1974)
Enewetak, 16-25 m 3,000-9,500   Huston, 1985 (Buddemeier et al., 1974)
Enewetak, >25 m 5,000-6,000   Huston, 1985 (Buddemeier et al., 1974)
P. lobata, Enewetak, 6-15 m 10,000-11,500   Huston, 1985 (Buddemeier et al., 1974)
Enewetak, 16-25 m 5,000-6,000   Huston, 1985 (Buddemeier et al., 1974)
Favia pallida
Enewetak, 0-5 m 5,500-7,500   Huston, 1985 (Highsmith, 1979)
Enewetak, 6-15 m 5,000-7,000   Huston, 1985 (Highsmith, 1979)
Enewetak, 15-25 m 4,000-7,000   Huston, 1985 (Highsmith, 1979)
Enewetak, 25-30 m 4,000-6,500   Huston, 1985 (Highsmith, 1979)
F. speciosa
Enewetak, 0-5 m 4,600   Huston, 1985 (Buddemeier et al., 1974)
Enewetak, 6-15 m 5,600-8,500   Huston, 1985 (Buddemeier et al., 1974)
Enewetak, 16-25 m 6,500-7,000   Huston, 1985 (Buddemeier et al., 1974)
Goniastrea retiformis
Enewetak, 0-5 m 8,000-10,000   Huston, 1985 (Buddemeier et al., 1974)
Enewetak, 6-15 m 5,000-9,500   Huston, 1985 (Highsmith, 1979)
Enewetak, 16-25 m 6,000   Huston, 1985 (Highsmith, 1979)
G. parvistella, 0-5 m 1,300-12,500   Huston, 1985 (Buddemeier et al., 1974)
Platygyra laminella, 6-15 m 6,700-8,000   Huston, 1985 (Buddemeier et al., 1974)
Reefs
Coral, atolls, windward 14,000   Kukal, 1971
Atolls, lagoon reefs, leeward 3,800   Kukal, 1971
Range, <5 m depth 1,100-20,000   Schlager, 1981
Range, 10-20 m depth 500-2,000   Schlager, 1981
Range, atolls, outer reefs 330-910   Kukal, 1971
Atlantic reefs
Alcaran Reef, Mexico, Acropora cervicornis reef 12,000 775 Macintyre et al., 1977
Alcaran Reef, Mexico, head-coral reefs 6,000 1,175 Macintyre et al., 1977
Grecian Rocks, Florida, windward 1,300 6,000 Shinn, 1980
Holocene, Miami, Florida 6,500 2,000 Lighty et al., 1978
Miami, Florida 740 4,900 Shinn, Hudson et al., 1977
Bal Harbor, Florida 380 6,300 Shinn, Hudson et al., 1977
Long Reef, Florida, windward 650 5,600 Shinn, Hudson et al., 1977
Carysfort Reef, Florida, windward 860-4,850 700-5,250 Shinn, Hudson et al., 1977
Big Pine Key, Florida, leeward 490-1,510 1,000-7,200 Shinn, Hudson et al., 1977
Dry Tortugas, Florida 1,910-4,470 130-6,000 Shinn, Hudson et al., 1977
St. Croix 15,200   Adey et al., 1977
St. Croix, algal ridge 6,000   Adey,1977
Hess Channel, St. Croix 2,300 3,400 Adey et al., 1977
Galeta Point, Panama, Acropora palmata reef
range 1,290-10,810 390-4,400 Macintyre & Glynn, 1976
average 3,900  
Galeta Point, reef-flat rubble, windward 600 5,500 Macintyre & Glynn, 1976
Boo Bee patch reef, Belize lagoon, leeward 1,600 8,800 Halley et al., 1977
Pacific reefs
Tarawa Atoll, windward 8,200 550 Marshall & Jacobson, 1985
Tarawa Atoll, windward 5,000-5,400 2,070 Marshall & Jacobson, 1985
Nishimezaki Reef, Ryukyus 3,900 7,400 Takahashi et al., 1988
Kikai-jima, Ryukyus 2,900-4,000 8,720 Takahashi et al., 1988 (Konishi et al., 1978)
Hanauma Reef, Hawaii 2,900 6,970 Takahashi et al., 1988 (Easton & Olson, 1976)
Great Barrier Reef, Holocene 200-600 9,000 Davies & Marshall, 1979
Great Barrier Reef, algal pavement 2,800 alkalinity meas. Davies & Marshall, 1979
Great Barrier Reef, reef, flat, coral zone, windward 3,100 alkalinity, 24 h Davies & Marshall, 1979
Great Barrier Reef, reef flat, windward 5,000 alkalinity, 24 h Davies & Marshall, 1979
Great Barrier Reef, leeward margin, leeward 6,000 alkalinity, 24 h Davies & Marshall, 1979
Carter Reef, Great Barrier Reef 260-2,200 320-1,940 Hopley, 1977
Orpheus Island, Great Barrier Reef 4,000 7,300 Takahashi et al., 1988 (Hopley & Bames, 1985)
Northern Great Barrier Reef 67-100 15 x 106 Davies, 1988
Central Great Barrier Reef 60-75 4 x 106 Davies, 1988
Southern Great Barrier Reef
(Heron Island, Wreck Reefs) 50-75 2-3 x 106 Davies, 1988
Enewetak, Holocene 320 13,000 Salter, 1984
Chiriqui Gulf, Panama, reef flat, average 2,640 2,825 Glynn & Mcintyre, 1977
Chiriqui Gulf, Panama, reef flat, range 1,100-4,800 5,585-210 Glynn & Mcintyre, 1977
Panama Bay, Panama, reef flat 1,300 4,150 Glynn & Mcintyre, 1977
Other carbonate environments
Andros Island, storm-tide layers 320,000-6,750,000 1.3-16.5 h Hardie & Ginsburg, 1977
Stromatolites 1,460,000 24 h Kukal, 1971
Bermuda stromatolites, subtidal 365,000 1-6 d Gebelein, 1969
Algal ridge, St. Croix, windward 6,000   Adey, 1977
Calcareous algae 2,000-7,000   Kukal, 1971
Calcareous algae, summer 12,000 1 m Kukal, 1971
Ooid shoals (range), windward 550-2,000 ≈3,000 Schlager, 1981
Great Bahama Bank, leeward 800-1,100 2,500 Schindel, 1980 (Cloud, 1962)
Great Bahama Bank, Andros lobe, leeward 200-850 ≈7,000 Enos, 1974
Little Bahama Bank 1,200 10,000 Sarg, 1988 (Hine et al., 1981)
Little Bahama Bank, Bight of Abaco (ave.), leeward 120 5,500 Neumann & Land, 1975
Little Bahama Bank, Bight of Abaco (core dates), leeward 200-300 ≈1,000 Neumann & Land, 1975
Florida, forereef slope 710 10,000 Enos,1977
Florida outer shelf margin, windward 490-1,010 6,000 Enos,1977
Florida inner shelf margin 180-610 5,000 Enos,1977
Florida inner shelf margin 220 10,200 Stockman et al., 1967
Rodriguez bank, Florida inner shelf margin 1,000 5,000 Wilson, 1975 (Turmel & Swanson, 1972)
Ceasar Creek Bank, Florida inner shelf margin 1,390 4,000 Warzeski, 1976
Florida Bay, mud banks, leeward 330 3,000 Stockman et al., 1967
Florida Bay, interbank, leeward 53 3,000 Stockman et al., 1967
Florida Bay, western mud bank, leeward 460 4,000 Schindel, 1980 (Scholl, 1966)
Florida Bay, western mud banks, leeward 620 4,000 Wanless & Tagett, 1989; K.K. Mukerji, unpub., 1987
Florida Bay, Cross Bank, leeward 1,585 1,700 E.A. Shinn & P.R. Rose, unpub., 1963
Florida Bay, Crane Key, leeward 1,000 3,000 Wilson, 1975 (Stockman et al., 1967)
Florida Bay, Cape Sable 4,000 50 Gebelein, 1977
Belize lagoon, adjacent patch reef 400-500 8,800 Halley et al., 1977
Galeta Point, Panama, forereef talus 2,500 2,500 Macintyre & Glynn, 1976
Galeta Point, fore-reef pavement, windward 600 3,750 Macintyre & Glynn, 1976
Galeta Point, back reef, windward 670 3,000 Macintyre & Glynn, 1976
Northeast Yucatan, lagoon, leeward 1,000 5,000 Wilson, 1975 (Brady, 1971)
Alcaran Reef, Mexico, coral rubble & sand, windward 2,500 4,500 Macintyre et al., 1977
Alcaran Reef, Mexico, reef flat, windward 2,000 3,500 Macintyre et al., 1977
Gulf of Mexico (oyster reef) 730 2,100 Schindel, 1980 (Shepard & Moore, 1955)
Persian Gulf 5-50 10,000 Schindel, 1980 (Samtheim, 1971)
Trucial Coast, Persian Gulf, lagoon, leeward 140 5,000 Schindel, 1980 (Kinsman, 1969)
Great Barrier Reef, sand flats 200 alkalinity meas. Davies & Marshall, 1979
Great Barrier Reef, reticulated lagoon, leeward 1,000 alkalinity, 24 h Davies & Marshall, 1979
Great Barrier Reef, deep lagoon 300 alkalinity, 24 h Davies & Marshall, 1979
Chiriqui Gulf, Panama, forereef slope, ave. 7,500 430 Glynn & Macintyre, 1977
Chiriqui Gulf, Panama, forereef slope, range 500-20,800 1,065-130 Glynn & Macintyre, 1977
Tidal flats, range 400-950 ≈3,000 Schlager, 1981
Cape Sable, FLorida, tidal flats 11,000 19 m Gebelein, 1977
Cape Sable, Florida, tidal flats 2,000-5,500 50 Gebelein, 1977
Andros Island, Bahamas, tidal flats 490 5,000 Hardie & Ginsburg, 1977
Andros Island, Bahamas, NW tidal flats 700 2,200 Wilson, 1975 (Shinn et al., 1965)
Andros Island, Bahamas, SW tidal flats 800 ≈5,000 Gebelein, 1975
Trucial Coast, Persian Gulf (intertidal) 400 5,000 Schindel, 1980 (Kinsman, 1969)
Sabkha, Trucial Coast, Persian Gulf 500 5,000 Wilson, 1975 (Kinsman, 1969)
Sabkha, Persian Gulf 29-94 3,000 Schindel, 1980 (Illing et al., 1965)
Sabkha Faishak, Persian Gulf 1,000 4,000 Wilson, 1975 (Illing et al., 1965)
Bathyal and abyssal deposits
North Atlantic, Holocene, average 88.5 11,000 Ericson et al., 1961
North Atlantic, Hologene, range 5-636 11,000 Ericson et al., 1961
North Atlantic, L. Pleistocene, glacial, average 63 50,000 Ericson et al., 1961
North Atlantic, L. Pleistocene, glacial, range 10→203 50,000 Ericson et al., 1961
California Borderlands, diatomaceous mud 880   Kukal, 1971
Yellow Sea, Sea of Japan, diatomaceous clay 50-200   Kukal, 1971
East India Sea, calcareous clay 850   Kukal, 1971
Atlantic and Pacific 10-150 ≈100,000 Schindel, 1980 (Ku et al., 1968)
North Pacific Ocean 11 7 x 106 Schindel, 1980 (Dymond, 1966)
Barbados (airborne) 0.6 9 m Schindel, 1980 (Delany et al., 1967)
Hemipelagic deposits
Continental slope, "blue mud", average 17.8   Kukal, 1971
Atlantic Ocean average 50-100   Kukal, 1971
Canadian slope, Atlantic 30-300 10,000 Lisitzin, 1972
European basin, Atlantic 485 11,000 Lisitzin, 1972
Senegal continental slope, Atlantic 200-418 11,000 Lisitzin, 1972
North American basin, west Atlantic 90-480 11,000 Lisitzin, 1972
Cuban slope, Caribbean 245-300 11,000 Lisitzin, 1972
Argentine Basin, S. Atlantic 3-45 11,000 Lisitzin, 1972
Demerara & Ceara Rise, W. equatorial Atlantic 3-400 15,000 Scholle et al., 1983 (Damuth, 1977)
Demerara & Ceara abyssal plain 15-40   Lisitzin, 1972
Ceara abyssal plain, equatorial Atlantic 200   Berger, 1974 (Hayes et al., 1972)
Nares abyssal plain (turbidites) 25-30 10,000 Kuijpers et al., 1987
W. Greater Antilles Outer Ridge, N. Atlantic 200-300   Kuijpers et al., 1987
E. Greater Antilles Outer Ridge, N. Atlantic 30   Kuijpers et al., 1987
Baleric abyssal plain, Mediterranean 160-520 20,000 Rupke,1975
North American rise, N. Atlantic 34-68 12,000 Schindel, 1980 (Emery et al., 1970)
North American abyssal plain, N. Atlantic 20 12,000 Schindel, 1980 (Emery et al., 1970)
Nares abyssal plain, N. Atlantic 2.5-10 10,000 Lisitzin, 1972
Sohm abyssal plain, N. Atlantic 20-360 10,000 Lisitzin, 1972
Cape Verde abyssal plain, mid-Atlantic 10-40 10,000 Lisitzin, 1972
Pernambuco abyssal plain, S. Atlantic 0.8 10,000 Lisitzin, 1972
Guinea abyssal plain, E. equatorial Atlantic 25-50 10,000 Lisitzin, 1972
Niger fan, E. equatorial Atlantic 25-65 10,000 Lisitzin, 1972
Angola abyssal plain, S. Atlantic 7.5-23 10,000 Lisitzin, 1972
Cape abyssal plain, S. Atlantic 1.5-12 10,000 Lisitzin, 1972
Bengal cone, Indian Ocean 64 10.2 x 106 Moore et al., 1974
North Pacific Ocean, blue mud 10   Kukal, 1971
Bering Straits, gray mud 80-4,500   Kukal, 1971
Bering Sea, basins 70-360 10,000 Lisitzin, 1972
California borderland 50-2,000   Berger, 1974 (Bandy, 1968)
California borderland 3-400 15,000 Scholle et al., 1983 (Prensky, 1973)
California borderland, ridges 50   Lisitzin, 1972 (Emery & Bray, 1962)
California borderland, proximal basin 1,800   Lisitzin, 1972 (Emery & Bray, 1962)
California borderland, distal basins 200-400   Lisitzin, 1972 (Emery & Bray, 1962)
California continental slope 80   Lisitzin, 1972
Kuril-Kamchatka trench 20-30 10,000 Lisitzin, 1972
Andean trench, E. Pacific 18-36 11,000 Schindel, 1980 (Lisitzin, 1972)
Antarctic slope, gray, silty clay 10-160   Kukal, 1971
Red clay
Oceanic average 7-13 12,000 Schindel, 1980 (Kuenen, 1950)
North Pacific Ocean 1-2   Opdyke & Foster, 1971
North Pacific (muddy) 10-15   Opdyke & Foster, 1971
North Pacific Ocean 0.2-6 10,000 Lisitzin, 1972
Tropical North Pacific 0-1   Berger, 1974
Indian Ocean 0.5-4.6 100,000 Schindel, 1980 (Kuznetsov, 1969)
North and South Pacific 2   Van Andel et al., 1975
Nares abyssal plain, N. Atlantic 5-10   Kuijpers et al., 1987
Brown clay
Oceanic average 2-10   Kukal, 1971
Nares abyssal plain, N. Atlantic 13-24 10-24.8 x 103 Kuijpers et al., 1987
Pelagic carbonate
Oceanic average, Globigerina ooze 10-80   Kukal, 1971
Oceanic average, Globigerina ooze 8-14 12,000 Schindel, 1980 (Kuenen, 1950)
Pacific Ocean, average 5.5 3 x 106 Davies & Worsley, 1981
Indian Ocean 10-40   Scholle et al., 1983 (Goldberg & Koide, 1963)
Indian Ocean average 11.9 3 x 106 Davies & Worsley, 1981
Atlantic Ocean, average 17.3 3 x 106 Davies & Worsley, 1981
Mid-Atlantic Ridge (crest) 1.7-185 11,000 Lisitzin, 1972
Mid-Atlantic Ocean 29 8,000 Schindel, 1980 (Nozaki et al., 1977)
North Atlantic Ocean 10-80   Scholle et al., 1983 (Ericson et al., 1961)
North Atlantic, Rockall Bank 25-75 10,000 Lisitzin, 1972
North Atlantic (40-50°N) 35-60   Berger, 1974 (Mcintyre et al., 1972)
North Atlantic (5-20°N) 14-40   Berger, 1974 (Schott, 1935)
Equatorial Atlantic Ocean 20-40   Berger, 1974 (Schott, 1935; Ericson et al., 1956)
South Atlantic Ocean 20-50   Scholle et al., 1983 (Ericson et al., 1961)
South Atlantic, Brazilian slope 30-50 10,000 Lisitzin, 1972
Caribbean Sea 24 116,000 Schindel, 1980 (Broecker & van Donk, 1970)
Caribbean Sea 12   Kukal, 1971
Caribbean Sea 20-110   Lisitzin, 1972
Caribbean Sea 10-60   Scholle et al., 1983 (Prell & Hay, 1976)
Caribbean Sea 28   Berger, 1974 (Emiliani, 1966)
Atlantic, Mediterranean 20-100   Kukal, 1971
Menorca Rise, Mediterranean, Quaternary 108 1.8 x 106 Hsü, Montadert et al., 1978
nanofossil marls
Black Sea, nannofossil ooze 100-300 3,000 Schindel, 1980 (Stoffers et al., 1978)
Mediterranean, pelagic ooze 50 10,000 Cita et al., 1978
Panama Basin, eastern equatorial Pacific 9-100   Scholle et al., 1983 (Swift, 1977)
Equatorial Pacific Ocean 10-25 10,000 Lisitzin, 1972
Western equatorial Pacific Ocean 11-50 106 Scholle et al., 1983 (Berger et al., 1978
Equatorial Pacific Ocean 5-18 106 Berger, 1974 (Hays et al., 1969)
Central equatorial Pacific Ocean 10-20   van Andel et al., 1975
Eastern equatorial Pacific Ocean 30   Berger, 1974 (Blackman, 1966)
East Pacific Rise (0-20°S) 20-40   Berger, 1974 (Blackman, 1966)
East Pacific Rise (30°S) 3-10   Berger, 1974 (Blackman, 1966)
East Pacific Rise (40-50°S) 10-60   Berger, 1974 (Blackman, 1966)
Northwest Providence Channel, Bahamas 15-22 183-93 x 102 Boardman and Neumann, 1984
Northwest Providence Channel, periplatforrn ooze 69-43 1,400-6,550 Boardman and Neumann, 1984
Biogenic siliceous sediments
Oceanic average, radiolarian ooze 5   Kukal, 1971
North & equatorial Atlantic Ocean 2-7   Berger, 1974 (Turekian, 1965)
South Atlantic Ocean 3-18 10,000 Lisitzin, 1972
South Atlantic Ocean 2-3   Berger, 1974 (Maxwell et al., 1970)
Pacific Ocean, diatom ooze 5-50   Kukal, 1971
Equatorial Pacific, siliceous ooze 4-5 106 van Andel et al., 1975
Equatorial Pacific, siliceous ooze 2-5   Berger, 1974
Equatorial Pacific Ocean 2-25 10,000 Lisitzin, 1972
Antarctic Ocean, radiolarian ooze 11-140   Scholle et al., 1983 (Hays, 1965)
Antarctic Ocean 0.7-32 10,000 Lisitzin, 1972
Indian Ocean, diatom ooze 5-20 100,000 Schindel, 1980 (Kuznetsov, 1969)
Gulf of California, varved diatomites 4,700-5,400   Lowe, 1976 (Calvert, 1966)
Vancouver Island fiord, varved diatomaceous sediment 4,000   Lowe, 1976 (Gross et al., 1963)
Freshwater diatomites (average) 300-1,000   Kukal, 1971
a. General format of the table is after Schindel (1980), as are many of the data. Other major sources are Kukal (1971), who does not generally indicate his sources, methods, or duration of observation, and Lisitzin (1972). The reference given in parentheses is the primary data source; those not listed among the references may be found from the secondary source cited.
b. Derived from many different types of measurement and from observations spanning vastly different time intervals. No corrections have been made for compaction. Longer periods ofobservation include some compaction, as well as more lacuna, than observations of shorter duration, a point emphasized by Schindel (1980). Rates are in Bubnoff (mm/103 yr = m/106 yr).
c. Time interval of observation is years, unless indicated otherwise.
*Thickness of individual flood deposits are reported as yearly rates on the assumption of annual flooding. These are more reasonable figures for modeling than calculated "instantaneous" sedimentation rates; moreover, the actual duration of flooding is rarely reported. It must be noted that most studies of floods are of exceptional events rather than of typical annual floods. For this reason some of the multiyear averages reported may be the most reasonable for simulations.

Accumulation rates

All-purpose sedimentation models incorporate accumulation rates for both terrigenous and carbonate sediments. These rates differ in general; they also respond in quite different fashions to changes in many other parameters. A wealth of data is available on rates of sediment accumulation (table 1). Well-documented values span such a range of rates and environments that the problem is in shopping: What are the appropriate values for the conditions to be simulated? The values in table 1 are grouped by depositional setting to facilitate selection. Different groupings, for example, by tectonic setting, may prove more appropriate for some models.

The time span of observation is an important determinant of sedimentation rates [Kukal, 1971; Schindel, 1980; see also Barrell (1917)]. In modeling this means that the time increment of the simulation may be important in choosing the appropriate sedimentation rate. Short-term observations invariably emphasize maximum rates produced by short-term events, such as floods or growth of organisms. It appears as though some rates approach infinity as the time of observation approaches zero (fig. 1). Clearly this is not the case, but the extrapolation emphasizes that short-term observations are not the most relevant to long-term considerations. The importance of duration of observation varies greatly among environments. Environments that suffer few perturbations (e.g., pelagic realms) have essentially uninterrupted sedimentation at rates that are virtually constant. In contrast, environments in which episodic events such as storms, floods, or turbidity currents dominate sedimentation typically have extreme short-term rates but intermittent deposition that modulates long-term averages. Schindel (1980) elegantly illustrated this phenomenon by plotting period of observation versus rate of sedimentation for a variety of environments. Figure 1 presents a series of such plots from the expanded data base of table 1.

Figure 1--Rates of sedimentation versus period of observation. The strong inverse relationship on most plots illustrates the gaps in the geologic record, the "long periods of boredom and short periods of terror" (Ager, 1981, p. 107). High "instantaneous" rates, for example, those generated by floods on floodplains and in deltas, are moderated by extended hiatuses. Note the contrast with more stable abyssal-plain environments. Even lakes show an inverse trend when ancient lacustrine environments are considered. Unfortunately, there seem to be no instantaneous rates on turbidity current deposition. Longer periods of observation include some apparent rate reduction because of compaction. No corrections for compaction have been made. Units of sedimentation rate are Bubnoffs (1 B = 1 mm/103 yr = 1 m/106 yr). Plots are logarithmic. Data from table 1. Expanded from Schindel (1980, fig. 1).

Eleven charts comparing rates of sedimentation vs. period of observation.

An important consideration incorporated in some two-dimensional programs [e.g., Lawrence et al. (1990)] is whether sediment is input from a point source, such as a river mouth, or a line source, such as a carbonate platform margin, or is uniformly distributed, as in pelagic sedimentation. A related consideration is that of throughput. As illustrated by a deltaic environment, some of the sediment input from the river accumulates in various subenvironments of the delta, but some is redistributed as hemipelagic or turbidite input to deeper environments. Menard (1961) provides some insight into apportionment to different environments in his analysis of diverse drainage basins (table 2).

Table 2--Partitioning of sediment by depositional environment.

Source Area Volume of Sediment
(106 km3)
Depositional Sites (% Volume)
Continental
& Shelf
Continental Rise Abyssal Plain
Appalachian Mountains 7.8 29 54 17
Mississippi drainage basin 11.1 81 11 8
Himalaya Mountains 8.5 49 1 50
Data from Menard (1961, p. 159).

In the carbonate realm sedimentation input from outside sources is typically minor or negligible compared with in situ production. The production rate is thus of prime importance. Carbonate production rates vary so greatly in magnitude and in response to various controlling factors that for some purposes it is desirable to model these responses. Lerche et al. (1987) explored the impact of some major controls on carbonate production rates and modeled their influence on the configuration of carbonate bodies. They introduced depth- and distance-dependent functions for food supply, light ("photosynthetically active radiation"), temperature, salinity, and oxygen concentration. These variables illustrate that, in general, climate can affect production rates, in addition to the character, of carbonate sediments profoundly. There are exceptions to the generalization that carbonate sediments are tropical (Teichert, 1958; Milliman, 1975, p. 204; Lees, 1975; Leonard et al., 1981; Rao, 1981), but data on these production rates are lacking. Carbonate production rates are generally greater in windward settings than in leeward ones, producing an inherent asymmetry in carbonate platforms. The rather sparse data to substantiate this (see table 1, Shallow-Water Carbonates) suggest that rates differ by factors of 2-4. Only the individual coral growth rates and carbonate fixation estimated from alkalinity measurements are truly production rates; other values are accumulation rates in a strict sense, although they may approximate production rates. It is normally more expedient to use accumulation rates; most of the data are in these terms (table 1, Shallow-Water Carbonates, Bathyal and Abyssal Deposits), and accumulation constitutes the sedimentary record (table 3).

Table 3--Sedimentation rates of "chemical" rocks in the geologic record.

Area Rate (B)b Period (yr)c Reference
Platform carbonates
Late Cambrian Whipple Cave Fm., Nevada 60 ≈9 Cook & Taylor, 1977
Late Cambrian tidal flats, Appalachians 25 18 Laporte, 1971
Late Cambrian subtidal, Appalachians 34 18 Laporte, 1971
Early Ordovician Ellenburger Group, Texas 15 27 Sarg, 1988 (Loucks & Anderson, 1980)
Early Ordovician Arbuckle Group, Oklahoma 110 27 after Wilson, 1975
Silurian pinnacle reefs, Michigan 13 14 Sarg, 1988 (Mesolella et al., 1974)
Late Silurian, Appalachians 100 ≈6 Laporte, 1971
Late Silurian, Midcontinent, USA 25 ≈6 Laporte, 1971
Early Devonian (Gedinnian) Helderberg Group, New York 15 7 Laporte, 1971
Middle Devonian Keg River platform, Alberta, Canada 14 11 Sarg, 1988 (Schmidt et al., 1980)
Late Devonian Swan Hills, Alberta, Canada 122 1 Sarg, 1988
Devonian (Givetian-Famenian), Canning basin 30 20 Schlager, 1981 (Playford & Lowrie, 1966)
Mississippian (Kinderhook-Meramec), Rocky Mountains 50-80 15 Schlager, 1981 (Rose, 1976)
Mississippian (Osage), Indiana 15 2 Brown et al., 1990
Mississippian (Osage), Indiana, tidal flats 350,000 1 x 10-6 Brown et al., 1990
Mississippian (Meramec-Chester), Rocky Mountains 100-150 8 Schlager, 1981 (Rose, 1976)
Pennsylvanian-Permian Nansen Fm. Sverdrup basin, Canada 37 52 Davies, 1977
Early Permian Wichita Fm., Texas 50 11 Sarg, 1988 (Silver & Todd, 1969)
Early Permian (Longyinian), Yangtze platform China 33-135 7 Enos, 1992
Early Permian (Qixian), Yangtze platform China 5-150 6 Enos,1992
Early Permian (Maokouan), Yangtze platform China 3-67 15 Enos, 1992
Late Permian (Longtan/Changxing), Yangtze platform, China 7-110 15 Enos, 1992
Permian Clear Fork Fm., Texas 365 1 Sarg, 1988 (Sarg & Lehmann, 1986)
Permian Grayburg Fm., Delaware basin, USA 160 1 Sarg, 1988 (Sarg & Lehmann, 1986)
Permian Capitan Fm., Delaware basin, USA 75 3 Schlager, 1981 (Harms, 1974)
Permian Capitan reef, Delaware basin, USA 55-83 9 Sarg, 1988 (Silver & Todd, 1969)
Permian San Andres Fm., Delaware basin 180 1 Sarg, 1988 (Sarg & Lehmann, 1986)
Triassic (late Anisian-Ladinian) Northern Calcareous Alps 100 7 Schlager, 1981 (Ott, 1967)
Triassic (Early Camian) Dolomites 300-500 4 Schlager, 1981
Late Triassic, Tethys (Alps, Apennines) 100   Bernoulli, 1972
Early Jurassic, Tethys (Alps, Apennines) 15-40 ≈20 Bernoulli, 1972
Jurassic Haynesville Fm., Texas 95 2 Sarg, 1988
Late Jurassic Smackover Fm., Arkansas 83 4 Sarg, 1988
Late Jurassic Friuli platform, southern Alps 30-45 20 Schlager, 1981 (Winterer & Bosellini, 1981)
Early Cretaceous Shuaiba, Middle East 155 1 Sarg, 1988
Mid-Cretaceous (Albanian-Cenomanian) Golden Lane, Mexico 100 15 Enos,1977
80 15 Wilson, 1975 (Coogan et al., 1972)
Cretaceous-Cenozoic, Andros well, Bahamas 35 120 Wilson, 1975 (Goodell & Garinan, 1969)
Cretaceous-Cenozoic, Sunniland field, Florida 30 120 Wilson, 1975
Mesozoic-Cenozoic, Persian Gulf (maximum) 30 200 Wilson, 1975
Late Eocene, Enewetak 170 3.4 Saller, 1984
Early Miocene, Enewetak 76 7.1 Saller, 1984
Late Miocene Terumbu Fm., S. China Sea 80-286 0.8-5.2 Sarg, 1988 (Rudolph & Lehmann, 1987)
Quaternary, Enewetak 11.5 0.59 Saller, 1984
Middle Miocene-Holocene, northern Great Barrier Reef 67-100 15   Davies, 1988
Pelagic and deep-water carbonates
Late Cambrian Hales Lst. (lower) Nevada 14 ≈9 Cook and Taylor, 1977
Late Cambrian Frederick Lst., Maryland 50 16 Reinhardt, 1977
Late Pennsylvanian-Early Permian Hare Fiord Fm. (lower), Sverdrup basin, Canada 16 28 Davies, 1977
Late Cretaceous Marne a Fucoidi (argillaceous) 15 ≈15 Bernoulli, 1972
Early Jurassic (Pliensbachian), High Atlas, Morocco 63-100 5-8   Evans & Kendall, 1977
Early Jurassic Monte Sant'Angelo Lst., Apennines (including platform debris) 13 20 Bernoulli, 1972
Early Jurassic, Comiola Fm., Apennines, & Sihiais Lst, Greece, (pelagic & calc. turbidites) 15-25 ≈20 Bernoulli, 1972
Middle Jurassic Lamellibranch Lst, Apennines, S. Alps, Greece 3-8 19   Bernoulli, 1972
Late Jurassic Oberalm Beds, Austrian Alps 17-51 5-15 Garrison & Fischer, 1969
Late Jurassic Cat Gap Fm., N. Atlantic 8-14 16 Jansa et al., 1979
Late Jurassic, Early Cret. Maiolica, Apennines, s. Alps 10 ≈23 Bernoulli, 1972
Late Cretaceous Chalk, UK
range 3-60 32 Scholle et al., 1983 (Hancock, 1975)
average 15 32
Late Cretaceous Chalk, Danish trough, North Sea 100   Scholle et al., 1983
Late Cretaceous, Tongue of the Ocean 8 ≈30 Bernoulli, 1972
Cretaceous, Italy
range 7-50   Scholle et al., 1983 (Arthur, 1979)
average 12  
Late Cretaceous chalks, Western Interior, USA
range 6.5-50   Scholle et al., 1983 (Kauffman, 1977)
average 35  
Early and middle Miocene, nannofossil marls, Menorca Rise, Mediterranean 103 7 Hsü, Montadert et al., 1978
Miocene Great Abaco Fm, N. Atlantic (intraclastic debris) 9-43 4.6-6.3 Jansa et al., 1979
DSDP cores, to site 335, 3 my averages 0.6-17 3 Davies and Worsley, 1981
Pacific Oceanic Plateaus
Aptian-Quaternary, Ontong Java Plateau 11.1 113 Jenkyns, 1978 (Moberly & Larsen, 1975)
Berriasian-Quaternary, Magellan Rise 8.9 131 Jenkyns, 1978 (Moberly & Larsen, 1975)
Barremian-Quaternary, Manihiki Plateau 7.8 116 Jenkyns, 1978 (Moberly & Larsen, 1975)
Berriasian-Quaternary, Shatsky Rise 4.9 131 Jenkyns, 1978 (Moberly & Larsen, 1975)
Cenomanian-Quaternary, Hess Rise 3.6 96 Jenkyns, 1978 (Moberly & Larsen, 1975)
Condensed Sequences
Late Devonian Cephalopodenkalk, Germany 1.5-2 14 Tucker, 1974
Late Devonian Griotte, France ≈7 7 Tucker, 1974
Late Triassic Hallstatt Lst., Austrian Alps 0.5-1.5 20 Garrison & Fischer, 1969
Early Jurassic Adnet Beds, Austrian Alps 0.6-1.0 15-25 Garrison & Fischer, 1969
Early-Middle Jurassic, Ammonitico Rosso, Apennines, s. Alps, Greece 2.5-6.5   Bernoulli, 1972
Early Pliocene nannofossil marls, Cretan Basin, Mediterranean 9 2.4 Hsü, Montadert et al., 1978
Pliocene nannofossil marls, Menorca Rise, Mediterranean 16 3.4 Hsü, Montadert et al., 1978
Pleistocene, Mediterranean 1 1.9 Cita et al., 1978
Siliceous rocks
Silurian-Mississippian or Devonian Caballos Novaculite, Texas 0.3-4.5 105-48 Folk & McBride, 1976
Devonian Arkansas Novaculite, Arkansas and Oklahoma (varved)
range 1,000-2,500 0.1 Lowe, 1976
average 1,250 0.1
Jurassic Radiolarite, Apennines 3-9   Schlager, 1974
Middle Jurassic Ruhpolding Radiolarite, Austria 0.7-1 20-30 Garrison & Fischer, 1969
Tithonian-Barremian Radiolarite Group, s. Alps 5.4 16 Bernoulli, 1972
Tithonian-Barremian Scisti ad Aptici, Apennines 5.8 16 Bernoulli, 1972
Tithonian-Barremian U. Posidonia Beds, Greece 3.1 16 Bernoulli, 1972
Eocene Bermuda Rise Fm, N. Atlantic, chert and siliceous mudstone 5-8 ≈10 Jansa et al., 1979
Miocene Monterey Fm, California, diatomite 8-200   Scholle et al., 1983 (Garrison & Douglas, 1981)
Miscellaneous pelagic rocks
Anhydrite, L. Permian Castile Fm, W. Texas-New Mexico 1,825 0.3 Dunham, 1972 (Udden, 1924)
Carbonaceous clays, Early Cretaceous Hatteras Fm., N. Atlantic 3-19 15-25 Jansa et al., 1979
Variegated clays, Late Cretaceous, N. Atlantic 1-3 27 Jansa et al., 1979
Hemipelagic mud, Eocene-Pleistocene Blake Ridge Fm., N. Atlantic 3-200 1.7-2.5 Jansa et al., 1979
Hemipelagic mud, Pleistocene, Nares Abyssal Plain, N. Atlantic 400-500 ≈0.05 Kuijpers et al., 1987
Evaporites
Late Silurian, Salina Group, Michigan basin 180 ≈6 Alling & Briggs, 1961
Late Silurian, Salina Group, Appalachians 150 ≈6 Alling & Briggs, 1961
Late Permian Castile Anhydrite, Texas-New Mexico (varved) 1,825 0.3 Dunham, 1972 (Udden, 1924)
Messinian, Sicily 160 1.2 Decima & Wezel, 1973
Messinian, DSDP Site 124, Baleric basin 67 1.2 Decima & Wezel, 1973
Messinian, DSDP Site 132, Tyrrhenian Sea 30 1.2 Decima & Wezel, 1973
a. Bubnoff, 1 B = 1 mm/103 yr = 1 m/106 yr.
b. Time intervals for stratigraphic units of Mesozoic and Cenozoic age are from Haq et al. (1987). Paleozoic intervals are from Palmer (1983).
c. Not all primary sources are listed in references; see secondary source for original reference.

Production rates exclude transported sediment, whereas this sediment is inherently incorporated in the accumulation rate. This difference leads to examination of the well-established dogma in carbonate sedimentology that most sediment is produced in situ and that lateral transport is minor or negligible [cf. Wilson (1975, p. 7)]. Scale must again be considered. Lateral transport may be appreciable on the scale of a reef, the nearest approximation of a point source in most carbonate realms. It is generally considered negligible on the scale of a basin or a platform, but it is clear that significant transport is necessary for lateral progradation of the slopes of carbonate platforms (Bosellini, 1984; Playford et al., 1989; Eberli and Ginsburg, 1989). Transport must likewise control sedimentation in tidal flats where in situ production is negligible. Periplatform carbonate ooze (Schlager and James, 1978), an important component of highstand accumulations in proximal portions of basins (Boardman and Neumann, 1984; Droxler and Schlager, 1985; Shinn, Steinen et al., 1989), demonstrates lateral transport of carbonates in suspension. The possibility of some lateral transport must therefore be considered to realistically model carbonate accumulation in two or three dimensions [cf. Spencer and Demicco (1989)].

One-dimensional models, focused on simulation of sequences by vertical accretion, generally ignore lateral transport (Read et al., 1986). This essentially denies the possibility of autocyclic sequences that are controlled by lateral progradation of sediment (Ginsburg, 1971). Current two-dimensional models generally deal with progradation in carbonates in essentially the same way as progradation in terrigenous clastics is treated (Demicco and Spencer, 1989; Lawrence et al., 1990; Bosence and Waltham, 1990). Accumulation produces vertical aggradation until the available space is filled; surplus sediment is then redistributed into adjoining areas.

More data exist for pelagic sedimentation rates in both carbonate and noncarbonate sediments than for any other environment, in part because of the Deep Sea Drilling Project, which calculates accumulation rates for each datable sedimentary interval. Time spans are typically a few million years. Only a reasonable sampling of these data is presented in tables 1 and 3. Impetus to systematically glean rates from the 100-plus volumes of the Deep Sea Drilling Project is reduced by the fact that pelagic sedimentation rates are generally the lowest encountered and the most stable. In some settings, however, basinal sedimentation rates may be of prime importance. Harris (1989) demonstrated the influence of basinal accumulation rates on progradation of Middle Triassic platform margins in the Dolomites of northern Italy. Progradation of platform margins is typically in response to increased shallow-water production rates or reduced accommodation space, but increases in basinal accumulation rates, especially through shifts to siliciclastics, volcanoclastics, or evaporites, also can dramatically increase progradation rates (Harris, 1989).

Mixed carbonate and terrigenous environments are not yet fully integrated into most models, even those capable of dealing with either terrigenous or carbonate sources [cf. Lawrence et al. (1990)]. Several new considerations are introduced. The cumulative sedimentation from both sources influences the overall sedimentation rate. When some threshold in terrigenous input is reached, carbonate productivity and therefore accumulation rate are apparently suppressed (Mount, 1984; Walker et al., 1983). Neither the threshold nor the rate of suppression can be quantitatively defined at present.

It is likely that the type of impinging terrigenous sediment must be considered in addition to its volume. Organisms can probably tolerate accumulation of sand and coarser sediments better than they can tolerate mud. Sand creates unstable, shifting substrates; it probably has little direct impact on the organisms' metabolism. Finer suspended sediment, however, has the more direct influence of fouling the feeding mechanisms of many carbonate-producing organisms or of smothering them (Ginsburg and Shinn, 1964; Wilson, 1975, pp. 1-3), although it may also produce fluid substrates inimical to epifauna. It is nevertheless probable that the tolerance of carbonate organisms to mud is higher than generally recognized. Many carbonate rocks include a high percentage of mud, carbonate or terrigenous, that was introduced over a long period of time. Some carbonate producers were excluded, but others survived or even flourished, and carbonate sedimentation continued [cf. Laporte (1969, p. 115)]. There is no indication that the inhibiting effects of terrigenous mud are any different from those of carbonate mud; carbonate-producing organisms cannot be expected to be mineralogists. Terrigenous mud influxes can, however, include land-derived excess nutrients that could further suppress carbonate productivity (Hallock and Schlager, 1986, p. 394). Productivity suppression from sediments or pollutants introduced by human activity (Weiss and Goddard, 1977; Smith et al., 1981) offers the best possibilities for quantification, an example of Nietzschean serendipity.

Lag time

Inundation of carbonate platforms during transgression apparently does not lead to the immediate onset of rapid production of carbonate sediment. Stated another way, carbonate production does not reach its full potential for a finite period (Schlager, 1981). Carbonate sediment accumulation therefore tends to lag the relative rate of sea-level rise, resulting in a deepening sequence (Read et al., 1986). The interval between initial inundation and onset of rapid sediment accumulation is the lag time. The formation of shoaling-upward platform cycles so common in the geologic record requires a lag time, according to current concepts of sedimentation (Read et al., 1986; Ginsburg, 1971). Otherwise, rapid carbonate sedimentation would maintain the sediment surface at sea level, and accumulation rates less than the relative rate of sea-level rise would form a continuously deepening sequence. Lag time is also essential to Ginsburg's (1971) autogenic cycles in which carbonate sediment builds up to sea level and progrades toward the platform edge, reducing the area of carbonate production until progradation ceases. To produce a transgression and begin a new cycle rather than maintain a steady-state aggradation, sediment accumulation must drop appreciably below the relative rate of subsidence for a finite period, the lag time.

Lag time has not been considered in terrigenous siliciclastic cycles because the capacity for the in situ production is lacking; sediment input does not necessarily change with submergence. It has been shown by analysis and by simulation that asymmetric shoaling-upward cycles can be produced by symmetric (e.g., sine wave) eustatic oscillations in sea level superposed on constant subsidence and sedimentation rates [cf. Jervey (1988)]. The rate of subsidence plus sea-level fall must exceed the rate of sediment supply near the inflection point of the sea-level curve, the point of maximum rate of fall. Such cycles could also be produced in carbonate sedimentation, of course, if the rate of sediment production were less than the combined maximum rates of sea-level fall and subsidence. Such solutions appear rather contrived, given the demonstrable rapid rates of carbonate production in shallow water (tables 1, Shallow-Water Carbonates; table 3, Platform Carbonates; Schlager, 1981). Moreover, the resulting cycles should invariably terminate with subaerial exposure of the upper part of the cycle and should commonly show a deepening portion of the cycle. Such elements are not rare in shoaling-upward platform cycles, but they do not appear to be the general case.

Although lag time has a profound effect on the character of simulated shoaling-upward cycles [cf. Read et al. (1986, p. 108), Goldhammer et al. (1987), and Koerschner and Read (1989)], processes that may cause sediment accumulation to temporarily lag subsidence are not recognized. It is clear that aggradation of sediment into the supratidal zone terminates carbonate production because of subaerial exposure. It is not clear why sediment production does not recommence immediately upon submergence. One possibility is that extremely shallow water results in temperature fluctuations or periodic exposure that inhibits carbonate production. The well-established increase in carbonate production rates with decreasing depth (fig. 2) may have an upper limit somewhat below sea level.

Figure 2--Carbonate sediment production as a function of depth. This curve is a hybrid, as were the original attempts to illustrate depth-related variations (Garrison and Fischer, 1969, fig. 22; Wilson, 1975, fig. 1-2); those curves have nevertheless proved fertile. This attempt to quantify the curve encountered a paucity of data on in situ production, especially at intermediate depths. These are needed to document the thresholds related to algal productivity, stressed by R. N. Ginsburg [cf. Wilson (1975) and Schlager (1981)], and the lower limit of coral-algal reef growth (≈70 m; James, 1977). In contrast, data on sedimentation rates, as opposed to production rates, in shallow-water and abyssal settings are abundant (table 1). Data used in constraining the curve are: BR, productivity of a Barbados reef in a sheltered setting (Steam et al., 1977). SR, compilation of accretion rates (mm/103 yr) of reefs at less than 5 m (15 ft) depths (Schlager, 1981). DR, compilation of accretion rates of reefs at 10-20 m (3065 ft) depth (Schlager, 1981). EF, EG, and EP, regressions of growth rate versus depth for three species of corals in Enewetak Atoll (Highsmith, 1979). JF and JB, regressions of growth rate versus depth of Acropora cervicornis on the forereef and backreef, respectively, in Jamaica (Tunnicliffe, 1983); these are linear growth rates of a branching coral, so they do not constrain accretion rates, but they should help define the variation with depth. C, T, and B, total productivity of carbonate mudbanks with varying degrees of restriction in south Florida [Upper Cross Bank, Tavernier Key, and Buchanan Banks, respectively (Bosence, 1988,1989)], paired with productivity from adjacent interbank areas (C', T', B'). PA, accumulation rates of pelagic carbonates from the equatorial Pacific; average for last million years and range for past 45 m.y. (van Andel et al., 1975). LY, the lysocline, a threshold of accelerated carbonate dissolution with depth, currently about 3,500-4,800 m (11,000-16,000 ft) (Scholle et al., 1983). CCD, the carbonate compensation depth, which ranges from <4,000 m (<13,000 ft) to >5,000 m (>16,000 ft) in the present oceans (Berger and Winterer, 1974). Benthic production is assumed to approach 0 with the disappearance of red algae at depths of approximately 250 m (800 ft). Pelagic production is a function of surface conditions, so it is shown constant with depth. Pelagic accumulation rates are affected primarily by dissolution on the seafloor. They decrease sharply below the lysocline and drop to 0 at the carbonate compensation depth. The conversion from production rates in grams per square centimeter per 1,000 years to sedimentation rates in Bubnoffs (mm/103 yr) varies with the porosity and mineralogy of the sediment; for example, 1 g/cm2/103 yr would equate to a sedimentation rate of 14.7 B for calcitic sediment with 75% porosity (mud), but only 8.5 B of aragonitic sediment with 40% porosity (sand). The conversion used, 10 B = 1 g/cm2/103 yr, would apply to aragonitic sediment with about 65% porosity.

Chart plots depth vs. accumulation rates.

Another possibility is that lag time has a physical basis that reflects lack of accumulation above some profile of equilibrium, such as the wave base (Enos, 1989). Accumulation of modern carbonate sediments in south Florida is confined below a threshold depth that appears to be a function of fetch, suggesting that fair-weather or storm wave base is the control. Critical depths are approximately 3 m (10 ft) in the open shelf margin of south Florida and 2 m (7 ft) in the protected inner shelf of Florida Bay. In the more open Atlantic setting of Antigua, West Indies, the threshold appears to be approximately 5 m (16 ft) (Weiss and Multer, 1988).

If either control, decreased productivity or wave base, is valid, the appropriate parameter is lag depth rather than lag time. The corresponding time is determined by the rate of change in relative sea level, which is a function of eustasy, subsidence, compaction, and accumulation rate. Goldhammer et al. (1987) used a constant lag depth of 1 m (3 ft) in simulations of shoaling-upward cycles in the Middle Triassic of the Dolomites. In contrast, Read et al. (1986) assigned various durations to lag time and observed how this influenced the cycles generated. It is obviously desirable that the lag parameter be an empirical input rather than an unknown. If the lag parameter is closely related to water depth, then the threshold will not be reached simultaneously across a sloping surface. Deeper areas would begin accumulating sediment while shoal areas still lie above the threshold depth. If wave energy is the ultimate control, then sheltered areas would have a shallower lag depth than more exposed areas.

In summary, appropriate lag depth or time cannot be satisfactorily specified at present. Specification will be possible only when the processes are better understood. If the threshold is energy related, it should become possible to make reasonable assumptions if enough is known about the regional setting. It would also be clear what parameters must be studied in modern environments to obtain better definition for modeling.

Accommodation space

Accommodation space (or accommodation potential) is the increment of room available for sediment accumulation. The upper limit of the space is the level above which net erosion will occur. In nearshore settings this is generally taken as sea level, although a profile of equilibrium with a basinward slope, a well-established concept in subaerial settings, probably also applies below sea level (Enos, 1977, pp. 106-107). The lower limit is the depositional interface, so the instantaneous accommodation space is approximately water depth. Because this increment varies with sediment accumulation, some arbitrary fixed datum, such as basement, is used to define accommodation space. Accommodation space then increases in response to eustatic rises in sea level, subsidence, and erosion. For some purposes these components can be lumped together in a single accommodation parameter. In general, it is essential to isolate the components and thereby illustrate their influence on accumulation patterns.

Sea-level fluctuations

The changes in sea level used for simulations are based on simple mathematical models or empirical sea-level curves. Some short-term simulations use a constant sea level or a uniform rate of change. This is realistic only for time spans less that 103-104 years; in fact, small-scale high-frequency oscillations may have durations of only a few hundred years (Dominquez et al., 1987). Sine or cosine functions of various amplitudes and wavelengths, which may be superposed on other functions, are also used (Bice, 1988; Jervey, 1988). Empirical sea-level functions include extrapolations of Quaternary sea-level curves (Watney et al., 1989), inferred from the coupling of glacial ice volume with δ18O content of pelagic foraminifers, reflecting sea-surface temperatures (Matthews, 1984). These data have been extended through the Cenozoic (Prentice and Matthews, 1988; Matthews, 1984). For longer-term simulations sea-level curves derived by sequence stratigraphy (Vail et al., 1977; Haq et al., 1987) have been used, although these also have their critics (Miall, 1986; Burton et al., 1987; Matthews, 1988; Christie-Blick et al., 1988; Gradstein et al., 1988).

Curiously, no one seems to have gone back to the roots and directly applied the complex wave functions resulting from the periodicities in orbital parameters (Milankovich, 1941) to pre-Pleistocene fluctuations in sea level. This would seem particularly appropriate in view of the current overwhelming acceptance of the Milankovich band of orbital fluctuations as a cause of short-term changes in sea level, as inferred from their postulated control on climate (Milankovich, 1941) and thereby on the waxing and waning of glaciers (Denton and Hughes, 1983).

Subsidence

Subsidence also has at least three components: tectonics, isostasy, and compaction. Some models simply assume a constant rate of subsidence at a given point, which translates into a linear boundary (constant slope) in two dimensions [cf. Jervey (1988)]. This simple model essentially incorporates all three components.

Tectonic subsidence

Commonly used functions for tectonic subsidence are based on crustal stretching and cooling (McKenzie, 1978; Steckler and Watts, 1978; Sclater et al., 1980; Watts et al., 1982). These functions accurately model the subsidence of passive continental margins as they move away from spreading centers. Subsidence rate depends on the spreading rate and time or distance from the spreading center. Such functions also have been applied to continental-interior and cratonic basins and to oceanic islands and trenches (Watts et al., 1982). A different model for the tectonic component is provided by the theory of flexure of an elastic plate (Turcotte and Schubert, 1982, pp. 125ff). This model, which predicts differential subsidence (including local uplift) across a basin, seems most appropriate to foreland basins subjected to rapid tectonic and sediment loading from adjacent mountain belts.

Differential warping is generally superposed on the stately submergence of passive margins, expressed as local arches and basins. The Cape Hatteras and Cape Fear arches and the Baltimore Canyon basin of the Atlantic margin of North America are examples. Warping, local faulting, and regional subsidence can be empirically adjusted from burial history, which gives temporal changes in depth to basement. On active margins these three factors are typically the dominant components of subsidence. This empirical approach incorporates a second component of subsidence, isostatic adjustment.

Isostatic subsidence

The crust subsides isostatically in response to sediment loading. The amount of subsidence can be approximated as the thickness of mantle material with a density of approximately 3.2 g/cm3 that would be displaced by an increment of unconsolidated sediment with the appropriate density. For short spans of time, <104 years, it may be necessary to consider the viscous delay in response to loading.

Isostatic response to loading by water is not normally considered in current modeling programs. It is unlikely to be important in areas that are submerged to depths greater than the increment of sea-level rise so that loading would be uniform. In shallow-water areas, however, water loading is differential and may become a significant parameter as resolution improves through more sophisticated modeling and improved parameter definition.

Compaction

The magnitude of compaction-induced subsidence can be three-quarters of the total thickness of muddy sediments because initial porosities are 70-80% (Hedberg, 1936; Rieke and Chilingarian, 1974; Enos and Sawatsky, 1981). Compaction-related subsidence can vary drastically with grain size and with degree of grain support, cementation, and sorting within sediments of the same composition. Variations related to composition are generally much less. Terrigenous muds are probably the most homogeneous sediment type in their compactional behavior; they generally show an exponential decline in porosity with depth. Absolute values vary appreciably among different basins, however (fig. 3). Far fewer empirical data are available on sandstones. The most comprehensive data set is that of G. I. Atwater and E. E. Miller (Blatt, 1979); these data indicate a linear rather than an exponential decrease in porosity with depth (fig. 4).

Figure 3--Representative fitted curves of porosity versus depth in mudrock. Locations: (1) Ciscaucasus, USSR. (2) Compilation, Tertiary and Quaternary. (3) Oklahoma, USA, Pennsylvanian and Permian. (4) Japan, Tertiary. (5) Venezuela, Tertiary. (6) Gulf Coast, USA, Tertiary. (7) Japan. (8) Compilation, mainly of locations 3 and 5. (9) Gulf Coast, USA, Tertiary. (10) Gulf Coast, USA, Tertiary. From Rieke and Chilingarian (1974, p. 42); references are given therein.

Porosity vs. depth for 10 locations.

Figure 4--Porosity versus depth of late Tertiary sandstones from Louisiana, US Gulf Coast. The 17,367 analyses have been averaged in thousand-foot intervals for calculation of the least-squares fit. Unfortunately, no error estimates or petrology are available. Unpublished data of G. I. Atwater and E. E. Miller (1965); from Blatt 979, p. 146).

Porosity vs. depth for Louisiana.

Compaction of carbonates has generated considerable controversy (Weller, 1959; Pray, 1960; Shinn, Halley et al., 1977; Shinn and Robbin, 1983). For pelagic carbonate muds a tremendous volume of data is available from the porosity versus depth curves generated by the Deep Sea Drilling Project. Available summaries are selective and out of date (fig. 5). The most comprehensive data set for shallow-water carbonates (fig. 6) is that of Schmoker and Halley (1982). Their results do not differentiate porosity loss by cementation from that by physical compaction. This point is important for modeling because changes in bulk volume, not in porosity, determine subsidence. Cement from external sources reduces porosity without reducing bulk volume; thus porosity curves converted to bulk volume loss may exaggerate compaction. This also applies to muddy carbonate rocks and to siliciclastics, although imported cement is probably less common. Conversely, secondary porosity produced by dissolution at depth (Schmidt and McDonald, 1979) might lead to underestimates of compaction. Reef carbonates are typically considered noncompactible in simulations and burial history routines, but even in these lithologies considerable compaction may ensue through pressure solution at depth [Mossop (1972) documents 13% compaction; Anderson and Franseen (1991) document 30%].

Figure 5--Porosity and depth in pelagic carbonates from the Pacific and Indian oceans. The left column is from Deep Sea Drilling Project (DSDP) site 167, Magellan Rise, central North Pacific. The right column has samples with 80% or more calcium carbonate from 11 other DSDP sites. Note that the ages apply only to site 167 (left-hand column). From Schlanger and Douglas (1974, p. 119).

Porosity vs. depth for Deep Sea Drilling Program site 167.

Figure 6--(a) Porosity versus depth from shallow-water carbonates, Pleistocene to Early Cretaceous in age, from south Florida. N = 1,302; 167 data points are gravity data from boreholes at shallow depths; the remainder are data points from porosity logs. From Schmoker and Halley (1982, p. 2,566). (b) Data from part a separated into dolomites (75-100% dolomite; N = 336) and limestones (75-100% CaCO3; N = 489). From Schmoker and Halley (1982, p. 2,569).

Porosity vs. depth for shallow-water carbonates.

Compaction curves for mixed terrigenous and carbonate sediments have not generally been isolated, although data are certainly available from the Deep Sea Drilling Project, where carbonate content is measured for the same intervals as porosity. Compaction curves of muddy terrigenous and carbonate sediments (figs. 3 and 5) generally do not differ sufficiently to preclude the use of the curve for the dominant sediment type or, preferably, a weighted average of the two curves for mixed sediments. A further consideration in the compaction of mixed sediments, not incorporated into existing models, is the possible effect on pressure solution. Conventional wisdom is that terrigenous clay in excess of approximately 10% enhances the susceptibility of carbonate sediments to pressure solution (Wanless, 1979; Bathurst, 1975; Dunnington, 1967). Documentation is inadequate and the mechanisms are not understood, but good empirical evidence that terrigenous clays do enhance pressure solution has been provided by McNeice (1987). It would be possible to incorporate this threshold in modeling, but there are scant quantitative data on pressure solution versus depth and even less on the relationship between pressure solution and terrigenous content of sediments.

In rapidly accumulating muddy sediments the possible generation of overpressure and the consequent retardation of compaction should be considered (Bradley, 1975; Plumley, 1980; Carstens and Dypvik, 1981; Shi and Wang, 1986). Overpressure resulting from sedimentation in excess of the rate at which pore fluids can escape would be amenable to modeling (Mudford and Best, 1989).

Erosion

Two environments of erosion must be considered in sedimentary modeling. Subaerial erosion contributes to sediment input and modifies the final configuration of the eroded area. Submarine erosion involves redistribution of sediment and corresponding changes in the final configuration.

Rates of subaerial erosion are complex functions of such factors as elevation, lithology, climate, and vegetation. The first two parameters may evolve from the simulation, and the last two, largely independent of parameters being modeled, are either ignored or indirectly specified by user-selected values. If the factors causing erosion are not a focus of the study, the typically complex interrelations of erosion are best treated by using empirical net rates of erosion. Estimates of erosion rates in various climates, terrains, and lithologies (tables 4-8) are provided by Corbel (1959a,b), Menard (1961), Ritter (1967), and Meybeck (1976). An empirical relationship potentially useful in modeling was derived by Ahnert (1970):

d = 0.1535h,     (1)

where d is the rate of denudation in Bubnoff units (mm/103 yr) and h is relief in meters. The relationship proposed by Schumm (1963) for relatively small drainage basins [< 1,500 mi2 (4,000 km2)] in semi-arid areas of the western United States may be more convenient in modeling:

log D = 26.9H - 1.7,     (2)

where D is the denudation rate in feet per 1,000 years (305 B) and H is the ratio of relief to the length of a drainage basin. The relief to length ratio is dimensionless and can be represented by the slope of a surface tilted above sea level in a simulation.

Table 4--Erosion rates by relief and climatea

Relief and Climateb Rate of Chemical
Erosion (B)
Total Rate
of Erosion (B)
Lowlands
Periglacial, permafrost (15/yr); Bear Lake, Canada 13 15
Continental, cold (28-75/yr); E. Canada, Sweden 17 19
Maritime, temperate (33/yr); NE Europe 24 ± 12 32 ± 11
Continental, temperate, Mississippi basin 15 59
Missouri basin (4.4/yr) 6 55
Mississippi basin (Ritter, 1967) 14 46.4
Mississippi basin, Quaternary (Menard, 1961)   42
Continental, and (0.7/yr); New Mexico 1 12
Tropical desert, central Sahara   1?
Hot, seasonally wet and dry; Paraguay 11 32
Tropical, humid (38/yr); Congo basin 15 22
Mountains
Periglacial (20/yr); Brooks Range, Alaska 12 300
Periglacial, humid (250/yr); Norway 325 580
Maritime, humid (482 cm); Juneau, Alaska 192 800
Maritime, cold-temperate (118/yr); glaciated Alps 99 ± 59 203 ± 181
Maritime, temperate; Swiss Alps (Blatt et al., 1980)   70-910
Alps, long-term (volume Rhone fan; Blatt et al., 1980)   400
Maritime, temperate; Mt. Ranier (Blatt et al., 1980)   3,000-8,000
Maritime, temperate; Appalachians (Menard, 1961)   8
Appalachians, southern (Blatt et al., 1980)   41
Long-term (Cenozoic detrital sed. vol.; Matthews, 1975)   27
Appalachians, northern (Blatt et al., 1980)   48
Long-term (Cenozoic detrital sed. vol.; Matthews, 1975   5
Mediterranean, high mountains (62/yr); Italy, France 78 449
Mediterranean, semi-arid (60/yr); Italy 40 100
Continental, temperate (90/yr); central Europe 50 ± 25 102 ± 61
Continental, Himalayas (Menard, 1961)   1,000
Himalayas, Kosi R. (100/yr)   1,145
Himalayas (Blatt et al., 1980)   720
Hot and humid (76/yr); Usumacinta, Mexico & Guatemala 30 92
Hot and and (4-10/yr); southeast USA 7 200
Hot and and (0.6/yr), Tunisia 8 130
After Corbel (1959b), except as noted. Corbel assumed rock densities of 2.5 g/cm3
to convert weights of material transported to surface lowering.
a. Units of erosion are Bubnoff (mm/103 yr = m/106 yr).
b. Figures in parentheses are runoff (precipitation minus evapotranspiration) in centimeters per year.

Table 5--Regional erosion rates in the United States

Region Mechanical (B) Chemical (B) Total Rate
of Erosion (B)
Period (yrs)
North Atlantic 23.7 18.5±2.2 42.2 4-10
South Atlantic & eastern Gulf of Mexico 16.8 15.9±5.3 32.7 4-8
Western Gulf of Mexico 34.6 8.1±5.3 42.7 6-9
Mississippi River basin 32.4 14.0±0.6 46.4 12
Colorado River basin 142.8 5.9±2.7 148.7 32
Pacific drainage in California 71.8 15.9±5.5 87.7 3-13
Columbia River basin 15.0 16.5±3.6 31.5 2-4
After Ritter (1967).
Suspended and dissolved loads only; bed loads not included. Assumed rock densities are 2.64 g/cm3
Erosion rates given in Bubnoff (mm/103 yr = m/106 yr).

Table 6--Rates of erosion from major drainages of the world

River Drainage Area
(103 km2)
Runoff
(cm/yr)
Mechanical
Erosion (B)
Chemical
Erosion (B)
Total
Erosion (B)
Amazon (Brazil) 6,300 88 39.9 17.6 47.5
Congo (Zaire) 4,000 31 5.0 4.4 9.4
Congo (Corbel, 1959b)*   38 6.3 14.4 20.7
Mississippi (USA) 3,267 17.8 36 15 51
Mississippi (Corbel, 1959b)*     42 14.2 56.2
Mississippi (Ritter, 1967)     32.4 14.0 46.4
Nile (Egypt) 3,000 2.8 14 2.2 16
Parana (Argentina) 2,800 20 15 7.6 23
Parana (Corbel, 1959b)*     20 10 30
Yenisey (Siberia, USSR) 2,600 21 1.9 11 13
Ob (Siberia, USSR) 2,500 15 2.4 7.6 10
Lena (Siberia, USSR) 2,430 21 2.4 14 16
Yangtze (PRC) 1,950 35 186   >186
Amur (Heilong) (Siberia) 1,850 19 5.2 4.1 9.3
Mackenzie (Canada) 1,800 17 25 15 39
Madeira (Brazil) 1,380 73 59 16 75
Hsi (Pearl) (PRC) 1,350 18 34   >34
Volga (Russia, USSR) 1,350 20 7.2 22 29
Zambesi (Mozambique) 1,340 17 28 4.4 33
Niger (Nigeria) 1,125 17 23 3.4 26
Murray (Australia) 1,070 2.1 11 3.1 14
St. Lawrence (Canada) 1,025 32.8 1.9 20 22
Orange (South Africa) 1,000 9.1 57 4.5 61
Ganges (Bangladesh) 975 37.8 203 30 233
Indus (Pakistan) 950 22 189 25 214
Orinoco (Venezuela) 950 100 34 20 54
Danube (Romania) 805 25.2 32 28 60
Mekong (Vietnam) 795 72.5 165 28 193
Negro (Brazil) 755 189 3.8 3.8 7.6
Huang (Yellow) (PRC) 745 10.7 814   >814
Columbia (USA) 670 37.5 16 20 36
Columbia (Ritter, 1967)     15.0 16.5 31.5
Kolyma (Siberia, USSR) 645 18.2 3.5   >3.5
Colorado (USA) 635 3.2 330 8.7 338
Colorado (Corbel, 1959b)*   4.4 207 10.7 217.8
Colorado (Ritter, 1967)     143 5.9 149
Chari (Chad) 600 6.9 2.5 1.7 4.1
Brahmaputra (Bangladesh) 580 104 519 49 568
Xingu (Brazil) 540 45 0.3 1.1 1.4
Tapajós (Brazil) 500 45 0.5 1.4 1.8
Dnieper (Ukraine, USSR) 500 10 0.8 8.3 9.2
Amu-Darya (Uzbekistan, USSR) 450 10 79 23 102
Irrawady (Burma) 430 98 265 - >265
Don (Russia, USSR) 420 6.6 5.2 13 18
Tigris-Euphrates (Shatt El Arab) (Iraq) 410 14 95 16 110
Maranon (Peru) 407 85 95 34 129
Ucayali (Peru) 400 76 116 52 169
Uruguay (Uruguay) 350 45 15 8.7 24
Magdalena (Colombia) 240 98 379 44 423
Rhine (Corbel, 1959b)* 225 49 1.9 28.3 30.2
After Meybeck (1976) with some additions and comparison. Estimates are based on suspended and dissolved loads only; bed loads are not included, except as noted by asterisks. Converted from weights assuming rock densities of 2.64 g/cm3 (Ritter, 1967). Units of erosion are Bubnoff (mm/103 yr = m/106 yr).
*Estimate includes bed load.

Table 7--Rates of erosion of rivers fed by glaciers

River Rate of
Erosion (B)
Hidden Glacier (Alaska; rapid advance) 30,000
Muir (Alaska) 5,000
Bosson (Chamonix, France) 1,800
Nant Blanc (French Alps) 1,600
Heilstuga (Norway) 1,400
Memurelven (Norway) 1,600
Auserfjötur (Iceland) 2,200
Jokullsá (Iceland) 2,200
Hoffelsjökull (Iceland) 3,200
Hofsjökull (Iceland) 1,800
Isortok (Greenland) 2,500
Saskatchewan (Canada) 2,000
From Corbel (1959b, p. 16). Includes estimates of bed load.
Assumed rock densities are 2.5 g/cm3. B = Bubnoff (mm/103 yr).

Table 8--Average erosion rates by continent.

Continent Mechanical Chemical Total
L/Ka G&Mb L/K G&M L/K G&M
North America 27.9 33.0 15.0 13.0 42.8 46.0
South America 35.3 23.2 20.9 11.6 56.2 34.8
Asia 62.8 122.4 16.2 12.6 79.0 134.9
Africa 17.7 6.2 9.6 9.0 27.3 15.2
Europe 16.5 9.8 11.9 18.0 28.4 27.8
Australia 12.2 10.0 4.2 1.0 16.4 10.9
World total 36.8 53.1 14.1 11.4 50.9 64.5
World total
(Ritter, 1967)
43-89 9.9 53-99
Units are Bubnoff: 1 B = 1 mm/103 yr = 1 m/106 yr. Estimates in tons per square kilometer per year were converted using 2.64 g/cm3 as the average density of crustal rocks (Ritter, 1967). Continental areas cited by Kukal (1971, p. 30) were used in calculations for uniformity between estimates.
a. Estimates from Lopatin (1952), in Kukal (1971, p. 30).
b. Estimates from Garrels and Mackenzie (1971, p. 120).

Erosion in subaerially exposed carbonate sediments is typically ignored. For short time spans (perhaps 104-105 yr) this is probably justified, because the initial stage of carbonate erosion is typically dissolution, which produces secondary porosity, not a lowering of the land surface. Over more extended periods, however, the collapse of larger pore spaces (e.g., caverns) can result in significant lowering of the landscape, which may be differential and partially predictable (Purdy, 1974). Most of the data on erosion in carbonate terrains are values for net transport of ions by rivers. With simplifying assumptions, these can be expressed as changes in elevation in the landscape (table 9).

Table 9--Rates of chemical erosion of limestones.

Climate/Location Runoff
(cm/yr)
Erosion
Rate (B)
Arctic, dry
Svalbard; Tanana, Alaska   40
Victoria Island, Canada   5
Somerset I, Canada (A & S) 10 2
Arctic, humid
Gold Creek, SE Alaska   530
Capilano, British Columbia   420
Svartisen, northern Norway   400
Maritime, cold
Vercors, French Alps   240
Lismore, Scotland   150
St. Casimir, Quebec   160
St. Théresé, Quebec   120
NW England (Sweeting)   40
Maritime, temperate
Derbyshire, England (A & S)   83
Fergus R., Ireland (A & S)   55
Mendip Hills, England (A & S) 82 63
Thames, England (Sweeting)   104
range   13-288
Lee, Essex, England (Sweeting)   63
range   23-155
Derwent, England (Sweeting)   66-197
Lesse, Belgium   27
Tamis, Yugoslavia   21
Alpine
Triglav, Yugoslavia (A & S) 280 130
Tolminka, Yugoslavia (A & S) 310 102
Tatry Mtns., Czechoslovakia (A & S) 122 50
range 110-160 33-95
Jura Mtns., Switzerland (A & S)   98
Continental, cold winters
Jasper, Alberta   40
Whitehorse, Yukon Terr., Canada   32
Fort Simpson, Mackenzie Terr., Canada   40
Gulf of Bothnia, Finland   30
Continental, temperate
Kentucky (Sweeting)   64
range   3-297
Coolamon, NSW, Australia (A & S) 120 24
Krakow, Poland 25 20
Texas, U.S.A. 4 5
Mediterranean
Postojna, Yugoslavia (A & S) 160 110
Bosna, Yugoslavia (A & S) 150 90
Trieste, Italy (A & S) 70 48
Senj, Yugoslavia (A & S) 46 28
Podovi, Yugoslavia (A & S) 25 15
Yugoslavia, karst mountains (humid)   60
Marseilles (dry)   10
South Algeria (arid)   6
Tropical, humid
Jamaica (A & S) 105 73
range 55-135 40-96
Puerto Rico (A & S) 70 41
Florida (A & S) 50 33
Kissimmee, Florida (Sweeting)   27
range   16-63
Usumacinta R., Mexico & Guatemala, mountains   45
Champoton R., Yucatan, lowland   16
Sources: Corbel (1959b, p. 19; 1959a), Sweeting (1964),
and Atkinson and Smith (A & S) (1976).
Units of erosion are Bubnoff: 1 B = 1 mm/103 yr= 1 m/106 yr.

The erosion of subaerially exposed terrigenous terrains has different implications for modeling than erosion of carbonate terrains in both the erosional and depositional realms. Carbonate material is removed primarily by dissolution and does not directly produce new sediment (Bosence and Waltham, 1990). Eroded terrigenous material must be redistributed and adjusted for any changes in bulk density between the original material and the unconsolidated sediment.

Submarine erosion by waves or by slumping and sliding can be significant. Some modeling programs test for slope stability after each increment of sediment is added and remove material where the simulated slopes are in excess of what is considered stable (Lawrence et al., 1990), a value commonly specified by the user. Marine engineering practice provides some empirical data (Roberts et al., 1980). Angles of large-scale slopes in carbonates and siliciclastics have been summarized by Schlager and Camber (1986); slopes of carbonate platforms tend to be much steeper and to covary with height (fig. 7). Few data are available on submarine erosion rates by waves and other shear forces. Rates of intertidal erosion of carbonates (table 10) (Trudgill, 1985) are apparently the nearest approximations available.

Table 10--Marine erosion: Coastal erosion of limestones

Location Substrate Rate
Red Sea Reef limestone 1,000
Barbados Beach rock (grazers) 1,000-2,000
Bikini Atoll, Marshalls Beachrock 300
Southwestern Australia Beachrock 27,000-67,000
Heron Island, Great Barrier Reef Beachrock 500
Aldabra Atoll, Indian Ocean Reef limestone (intertidal) 500-4,000
Aldabra Atoll, Indian Ocean Reef limestone (subaerial) 260
Aldabra Atoll, Indian Ocean Reef limestone (with sand) 1,250
Aldabra Atoll, Indian Ocean Reef limestone (no sand) 1,010
From Trudgill (1985, p. 159). Surface lowering in B (mm/103 yr)

Figure 7--Slope angles on large slopes: angle of upper one-third of slope versus height of slope. Contours indicate concentrations of 0.5%, 1%, and 2% of total sample in unit area of 0.25 km x 0.05 tan S, measured as the moving average of 9 unit cells. Carbonate sample includes Bahamas and Marshall Islands (atolls) (N= 413). Siliciclastic sample is based on Atlantic continental slopes (N = 72). Carbonate slopes steepen with height up to at least 5,000 m (15,000 ft). Siliciclastics follow this trend only to 500 m (1,500 ft); slope height above 500 m has no influence on steepness of siliciclastic slopes. Data of Schlager and Camber (1986); from Schlager (1988).

Slopes steepen with height for carbonates; after a certain steepness height does not affect slope for siliciclastics.

Summary and conclusions

This review was undertaken to summarize optimum sources of input data for sedimentation simulations and to point out where additional data or refinement of concepts is needed from fieldwork. The trinity of parameters-accumulation rates, lag time, and accommodation space-are judged to be the most fundamental. Available data are summarized in the figures and tables. Urgent needs are an elucidation of lag time, quantification of compaction by pressure solution, and evaluation of the effects of siliciclastic influx on both carbonate sedimentation and pressure solution.

Acknowledgments

Lynn Watney persuaded me to undertake this summary, and Evan Franseen shepherded the later stages of preparation. Hal Wanless made many perceptive criticisms that helped to improve the manuscript, and Tony Simo smoothed some rough edges. Randy Farr was a skilled and patient coach at computer drafting. Special thanks are due to Nuria Wells for processing countless drafts, including the horrendous table 1.

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