OCS Study

BOEM 2020-004

U.S. Department of the Interior

Bureau of Ocean Energy Management

Gulf of Mexico OCS Region

Cooperative Agreement

Coastal Marine Institute

Louisiana State University

Coastal Marine Institute

Response of Later Quaternary Valley

Systems to Holocene Sea Level Rise

on the Continental Shelf Offshore

Louisiana: Preservation Potential of

Paleolandscapes
OCS Study

BOEM 2020-004

Published by

U.S. Department of the Interior

Bureau of Ocean Energy Management

Gulf of Mexico OCS Region

Cooperative Agreement

Coastal Marine Institute

Louisiana State University

Coastal Marine Institute

Response of Later Quaternary Valley

Systems to Holocene Sea Level Rise

on the Continental Shelf Offshore

Louisiana: Preservation Potential of 

Paleolandscapes

Authors

Paul V. Heinrich

Michael Miner

Robert Paulsell

Richard P. McCulloh

January 2020

Prepared under BOEM Cooperative Agreement

M12AC000020 by

Louisiana Geological Survey

Energy Center

Louisiana State University

Baton Rouge, Louisiana 70803
DISCLAIMER

Study collaboration and funding were provided by the US Department of the Interior, Bureau of Ocean 

Energy Management (BOEM), Environmental Studies Program, Washington, DC, under Agreement 

Number M12AC000020. This report has been technically reviewed by BOEM, and it has been approved 

for publication. The views and conclusions contained in this document are those of the authors and should 

not be interpreted as representing the opinions or policies of the US Government, nor does mention of 

trade names or commercial products constitute endorsement or recommendation for use.

REPORT AVAILABILITY

To download a PDF file of this report, go to the US Department of the Interior, Bureau of Ocean Energy 

Management Data and Information Systems webpage (http://www.boem.gov/Environmental-StudiesEnvData/), click on the link for the Environmental Studies Program Information System (ESPIS), and 

search on 2020-004. The report is also available at the National Technical Reports Library at 

https://ntrl.ntis.gov/NTRL/.

CITATION

Heinrich, PV, Miner M, Paulsell R, McCulloh RP (Louisiana State University, Baton Rouge). 2020. 

Response of Late Quaternary Valley systems to Holocene sea level rise on continental shelf offshore 

Louisiana: preservation potential of paleolandscapes. 104 p. New Orleans (LA): US Department of 

the Interior, Bureau of Ocean Energy Management. Report No.: BOEM 2020-004. Agreement No.: 

M12AC000020.
ii

Contents

1. Introduction .............................................................................................................................................6

1.1 Project Overview and Objectives .........................................................................................................................9

1.2 BOEM Information Needs Addressed................................................................................................................10

2 Methodology...........................................................................................................................................11

2.1 Background and Archival research ....................................................................................................................11

2.2 Literature Review and Synthesis .......................................................................................................................12

2.3 Offshore Lease Block Hazard Survey Reports Database Development ............................................................12

3. Northern Gulf of Mexico Geology and Sea Level History.................................................................15

3.1 Geologic Setting ................................................................................................................................................15

3.2 Sea Level Change: An Overview .......................................................................................................................17

3.2.1 Eustatic controls on sea level.....................................................................................................................17

3.2.2 Vertical displacement of oceanic and continental crust ..............................................................................18

3.2.3 Halokinesis and surface deformation .........................................................................................................20

3.2.4 Sediment compaction and subsidence.......................................................................................................20

3.3 Relative Sea Level History.................................................................................................................................20

3.3.1 MIS 5, Last Interglacial...............................................................................................................................21

3.3.2 Last Interglacial highstands, MIS5, Gulf of Mexico.....................................................................................21

3.3.3 Falling-stage sea level: MIS 4 and 3 ..........................................................................................................24

3.3.4 Last Glacial Maximum sea level MIS 2–LGM.............................................................................................27

3.3.5 Holocene Sea Level: MIS 1........................................................................................................................28

4. Shelf Morpho-sedimentary Response to Sea Level Change............................................................30

4.1 Transgressive Surfaces .....................................................................................................................................30

4.1.1 Coastline flooding (ravinement) surfaces ...................................................................................................30

4.1.2 Shoreface ravinement surface ...................................................................................................................31

4.1.3 Tidal ravinement surfaces ..........................................................................................................................33

4.2 Paleovalleys and Channel Belts ........................................................................................................................34

4.2.1 Forced regression, paleovalleys, and channel belts...................................................................................34

4.2.2 Transgressions, paleovalleys, and channel belts .......................................................................................35

4.3 Deltas and Chenier Plain ...................................................................................................................................36

4.3.1 Regressive and transgressive phases of deltas .........................................................................................36

4.3.2 Regressive and transgressive phases of Louisiana Chenier Plain.............................................................37

4.3.3 Forced regression, deltas, and beach ridges .............................................................................................37

5 Preservation of Paleolandscapes and Cultural Resources on the Northern Gulf of Mexico 

Continental Shelf.......................................................................................................................................39

5.1 Landscape: Cultural Resources Associations....................................................................................................39

5.2. Paleolandscape and Cultural Resources Preservation.....................................................................................42

5.2.1 Transgression: coastline flooding surfaces related to paleolandscape preservation..................................42

5.2.2 Transgression–shoreface ravinement ........................................................................................................43

5.2.3 Tidal ravinement.........................................................................................................................................44

5.2.4 Normal regression–progradation................................................................................................................45

5.2.5 Forced regression ......................................................................................................................................49

6. Stratigraphy of Study Area ..................................................................................................................50
iii

6.1 Allostratigraphic Approach to Pleistocene Unit Definitions.................................................................................51

6.1.2 Beaumont Alloformation (Prairie Allogroup) ...............................................................................................52

6.1.3 Avoyelles Alloformation ..............................................................................................................................58

6.2 Trinity-Sabine Unconformity...............................................................................................................................60

6.3 Deweyville Allogroup..........................................................................................................................................61

6.4 The Holocene-Pleistocene Surface....................................................................................................................65

6.5 Mermentau Alloformation...................................................................................................................................66

6.6 Unnamed Ravinement Surface..........................................................................................................................69

6.7 Sabine Bank Alloformation.................................................................................................................................69

7. Paleogeography ....................................................................................................................................70

7.1 Paleochannels and Their Fills............................................................................................................................70

7.2 Channel Belts and Deposits...............................................................................................................................76

7.3 Paleovalleys and Their Fills ...............................................................................................................................76

7.4 Paleochannels Compared to Channel Celts and Paleovalleys ..........................................................................78

7.5 Southwest Outer Continental Shelf Study Area .................................................................................................79

7.6 Paleoshorelines and Inner Shelf Shoals ............................................................................................................82

8. Synthesis and Summary ......................................................................................................................83

8.1 Site Preservation and Stratigraphy ....................................................................................................................84

8.2 Recommendations for Management and Protection..........................................................................................85

Works Cited ...............................................................................................................................................87
iv

List of Figures

Figure 1.1. Late Pleistocene eustatic sea level curve constructed from Pacific Gas and Electric 

Company (2011)..........................................................................................................................7

Figure 1.2. Eustatic sea level curve for the last glacial-interglacial cycle according to, compiled, and 

modified from Lambeck and Chappell (2001), Chappell (2002), Lambeck et al. (2002), 

Cutler (2003), Peltier (2004), and Pacific Gas and Electric Company (2011). ...........................8

Figure 2.1. Southwest Louisiana Outer Continental Shelf Study Area........................................................11

Figure 2.2. Data Distribution Southwest Louisiana Outer Continental Shelf Study Area............................13

Figure 2.3. Distribution of paleolandforms in southwest Louisiana Outer Continental Shelf Study Area, 

as mapped by OSC Hazard studies and Pearson et al. (1986)................................................14

Figure 3.1. Diagrammatic interpretation of Quaternary depositional systems of the Texas Coastal 

Plain. .........................................................................................................................................16

Figure 3.2. Geomorphic map of Southwest Louisiana Pleistocene paleochannels. ...................................22

Figure 3.3 MIS 5–3 channel belts of the western Louisiana continental shelf, as modified from Suter 

and Berryhill (1985) by Anderson et al. (2016). ........................................................................25

Figure 3.4. Paleogeographic map showing major depositional systems that existed on the shelf 120 ka 

to 22 ka (stages 5e–3 highstand)..............................................................................................26

Figure 3.5. Paleogeographic map showing the major depositional systems that existed on the shelf 

and upper slope during the 22 ka to 16 ka (MIS 2) lowstand. ..................................................28

Figure 4.1. Marine transgression by shoreface retreat in which shoreface erosion associated with a 

rise in sea level removes older coastal deposits, except within the paleovalley, and 

lagoonal flooding surface and creates marine ravinement surface. .........................................31

Figure 4.2. Marine transgression by in-place drowning during period of rapid relative sea level rise.........32

Figure 4.3 Normal regression in which coastline progrades gulfward in response to sediment 

accumulation during period of constant or slowly rising relative sea level. ..............................36

Figure 4.4. Forced regression in which coastline moves gulfward in response to relative sea level fall. ...38

Figure 5.1. Preferred prehistoric site locations in the Mississippi Deltaic Plain...........................................39

Figure 5.2. Cross section of terrace of “High” Deweyville at McNeill Ranch Site, Victoria County, 

Texas.........................................................................................................................................40

Figure 5.3. Location of shell middens and shell content in a hypothetical coastal estuary.........................41

Figure 5.4. Preservation potential of cultural deposits within a major paleovalley fill (e.g., Sabine) in 

respect to depth of erosion of the ravinement surface..............................................................44

Figure 5.5. Distribution of initial occupation sites in a prograding beach sequence. From A-oldest; to Eyoungest....................................................................................................................................46

Figure 5.6. Shell mound on a subsided natural levee ridge enveloped by marsh mud...............................47

Figure 5.7. Schematic dip cross section, central coastal plain, Texas........................................................48

Figure 6.1. Regional dip cross section showing relationships between outcropping formations and 

subsurface stratigraphy, northwestern Gulf of Mexico Coastal Plain. ......................................51

Figure 6.2. Regional dip cross section showing relationships between stratigraphic units within the 

study area and onshore Southwest Louisiana..........................................................................53

Figure 6.3. Correlation chart of regional Late Neogene strata showing the relationship between 

biostratigraphy, coastal onlap curve, and marine oxygen isotope stages. ...............................55
v

Figure 6.4. Late Pleistocene stratigraphy of Southwest Louisiana and offshore study region showing 

the relationship between major coastal plain stratigraphic units, global sea level, 

paleomagnetic chronology, and chronostratigraphy. ................................................................57

Figure 6.5. Cross-section across the Trinity-Sabine Paleovalley. ...............................................................61

Figure 6.6. Holocene–Pleistocene Surface, 5-m Contour Interval. .............................................................62

Figure 6.7 North-South cross-section across the Louisiana Chenier Plain showing sedimentary facies 

and bounding unconformities of the Mermentau Alloformation. ...............................................67

Figure 7.1. Nomenclature for Fluvial Valley Landforms...............................................................................71

Figure 7.2. LiDAR DEM of Amite River Valley and Paleovalley (PV) showing channels, paleochannel, 

and intervening interfluves. .......................................................................................................72

Figure 7.3. Hierarchy of fluvial architectural elements.................................................................................74

Figure 7.4. Peveto Channel sand body on the OCS used for Holly Beach Restoration interpreted as 

Late Pleistocene channel belt. ..................................................................................................75

Figure 7.5. Excerpt from plot of fluvial features mapped by offshore hazard surveys. ...............................79

Figure 7.6. Excerpt from plot of fluvial features mapped by offshore hazard surveys and Suter (1987). ...81

Figure 7.7. Excerpt from plot of fluvial features mapped by offshore hazard surveys and Suter (1987). ...82
6

1. Introduction

Understanding the complex response of coastal landforms to sea level rise is important for evaluating 

preservation potential of paleolandscapes and identifying sand resources on the Outer Continental Shelf 

(OCS). During the last glacial maximum, which ended approximately 18,000 years before present, sea 

level in the northern Gulf of Mexico was approximately 120 m lower than it is today and the lowstand 

shoreline was positioned approximately coincident with the shelf edge. As a result, intrabasinal coastal 

streams extended across the shelf within deep incised valleys separated by interfluves. As sea level rose, 

the incised valleys were infilled and coastal systems migrated landward across the shelf. This valley 

infilling consisted of both aggrading fluvial deposits and backstepping (migrating landward with sea-level 

rise) estuarine systems that developed within the flooded river valleys. As sea level rose, the retreat of 

coastal and bay shorefaces ultimately encountered and significantly eroded valley fill deposits and 

flanking interfluve surfaces. This produced regionally persistent ravinement surfaces. Also, shoreface 

retreat and erosion significantly affected the surfaces and sediments of any attendant deltaic systems. 

Typically, the complex erosional and depositional responses to sea level rise associated with shoreface 

retreat resulted in all but the basal portions of coastal and estuarine deposits being truncated and reworked 

by wave and tidal currents. There is little to no preservation potential of coastal lithosomes seaward of the 

modern shoreface. However, there may be unique situations associated with fluvial aggradation and

subsequent estuarine deposition within paleovalleys that could potentially provide an opportunity for 

preservation of buried geomorphic surfaces and intact prehistoric landscapes. Preserved fluvial and 

estuarine sediments as part of paleovalley fills on the OCS is important because they present rare 

opportunities for preservation of paleolandscapes and associated cultural sites that survived sea level rise 

because they accumulated at elevations below the depths of shoreface ravinement. Moreover, the fluvial 

sediments underlying these terraces are inferred to be locally sand-prone channel bed and point bar 

deposits, a potential sand source for coastal restoration projects.

Since the mid-Pleistocene (approximately 900 ka), the periodic growth of continental ice sheets has 

reduced the volume of seawater in the ocean basins. These reductions in the volume and thermal 

contraction of seawater have repeatedly lowered global, eustatic sea levels on the order of 120 m below 

present levels during periods of maximum glaciation (Chappell et al. 1996) (Figure 1.1). Due to the effect 

of glacial isostatic adjustment, relative sea level was lowered on the order of about 90 m below the 

present sea level (Bart and Ghoshal 2003, Roberts et al. 2004). During the last glacioeustatic cycle, 

Marine Isotope Stages (MIS) 5 to 1, sea level fell and rose within the Gulf of Mexico as ice sheets in 

North and South hemispheres expanded and shrank. At the beginning of the last glacioeustatic cycle about 

120 ka (MIS 5), sea level in the Gulf of Mexico was at a maximum highstand of ~6 m above present. 

Subsequently, it episodically fell until the middle of MIS 4, at about 70 ka when it was approximately 80 

m below present sea level. Then sea level rose to about 15 m below present sea level during the early part 

of MIS 3. Following this intermediate highstand, sea level again episodically fell to about 120m below 

present by about 22 ka during MIS 2. This lowstand lasted from about 22 to 16 ka. Between 16 and 4 ka, 

sea level rose over 100m as continental ice sheets melted (Anderson et al. 1992 and 2004, Blum and Price 

1998) (Figure 1.2).
7

Figure 1.1. Late Pleistocene eustatic sea level curve constructed from Pacific Gas and Electric Company (2011). 

Shackleton (2000)’s curve is based on marine oxygen isotope record corrected using Vostok ice core oxygen isotope record. The curve of 

Waelbroeck et al. (2002) is based on regression analysis of relative sea-level data from coral reef and oxygen isotope data.
8

Figure 1.2. Eustatic sea level curve for the last glacial-interglacial cycle according to, compiled, and modified from Lambeck and 

Chappell (2001), Chappell (2002), Lambeck et al. (2002), Cutler (2003), Peltier (2004), and Pacific Gas and Electric Company 

(2011).
9

Within the Gulf of Mexico, these sea-level fluctuations have resulted in drastic changes in the position of 

the shoreline. During the peak of the MIS 5e highstand about 120 ka the Gulf of Mexico shoreline was 

several kilometers inland of the present shoreline along the Ingleside relict shoreline trend in Texas and 

Southwest Louisiana. In Southeast Louisiana, its location is uncertain, although during MIS 5a the Gulf 

of Mexico’s shoreline lay along the Ponchatoula strandplain. Following MIS 5, the alternating lowstands 

and highstands have resulted in the development of entrenched stream valleys and deposition of shelfphase deltas on the Louisiana Continental Shelf. These and lesser fluctuations of base level also have 

resulted in well-defined unconformity-bounded, allostratigraphic, packages of fluvial sediments and their 

associated terrace surfaces within these entrenched valleys (Anderson et al. 1992 and 2004, Blum and 

Price 1998, Blum and Törnqvist 2000).

Although their ages of deposition predate the established timing of prehistoric human occupation in North 

America, many of the terraces within coastal plain valleys and paleovalleys were subaerial landscapes 

open to prehistoric occupation during MIS 2. In some cases, fluvial, estuarine, and eventually marine 

deposits buried these paleolandscapes and associated floodplains as sea level rose between 16 and 4 ka. 

Lying within entrenched valleys beneath shoreface erosion and the resulting ravinement surface, these 

surfaces and any associated archaeological deposits might have survived the marine transgression 

associated with postglacial sea-level rise (Blum and Price 1998, Gagliano et al. 1982, Pearson et al. 1986,

Stright 1990 and 1995).

In addition, the Pleistocene fluvial packages that underlie or are laterally bounded by these terrace

deposits are a potential sand resource for long-term coastal restoration and beach nourishment programs. 

The fluvial sediments underlying these terraces are inferred to be locally sand-prone channel bed and 

point bar deposits. Therefore, some potential sources of sand from the Louisiana-Texas Continental Shelf 

are also areas that potentially contain preserved, buried paleolandscapes. Thus, it is important for 

archaeologists studying the continental shelf to understand the stratigraphy, geology, geochronology, and 

geoarchaeology of the deposits filling the incised valleys and the attendant erosional surfaces pervasive 

on the Louisiana-Texas Continental Shelf (Gagliano et al. 1982, Pearson et al. 1986, Stright 1990 and

1995). 

1.1 Project Overview and Objectives

The objective of the study was to develop a conceptual model for Late Pleistocene to recent modification 

of the northern Gulf of Mexico coastal plain during the Holocene transgression. The purpose of this 

model is to evaluate the preservation potential of coastal plain paleosurfaces, e.g., uplands, channel belts, 

floodplains, terraces, and barrier islands as sea level submerged the Late Pleistocene coastal plain and 

transformed it into the modern continental shelf. Eventually, it should be possible to map preserved

paleosurfaces and evaluate the potential existence of prehistoric cultural resources on the OCS. In 

addition, the same conceptual model can be used to evaluate the potential existence of sand resources on 

the OCS suitable for restoration projects and quantitatively predict their distribution. Essentially the same 

geologic factors that govern the preservation or destruction of paleosurfaces within the OCS also 

influence distribution of sand within shelf-coastal systems. Finally, this study developed a basic 

stratigraphic framework model that can be used to inform future geologic and high resolution geophysical 

data collection to address questions about paleosurfaces and sand resources. 
10

1.2 BOEM Information Needs Addressed 

BOEM is tasked with making available OCS resources for expeditious and orderly development in a 

manner that safeguards environmental and cultural resources. A number of uses of Gulf of Mexico 

resources, such as fishing, energy development, recreation, sand dredging, affect underwater cultural 

resources, such as historic shipwrecks and submerged prehistoric sites. However, little is known about

early human occupation on the continental shelf and the potential for site preservation. Sedimentological 

and geoarchaeological studies have shown that deposits on continental shelves respond to modern 

hydrological processes. Moreover, during marine transgression significant erosion by waves and tides as 

the coastal system migrates up the shelf results in significant erosion of the shelf. This observation 

renders invalid the concept that the continental shelves consist of an unmodified, drowned coastal plain. 

However, sedimentary deposits and subaerial paleosurfaces, with which intact and potential significant 

cultural deposits might be associated, can be preserved under predictable circumstances (Coastal 

Environments Inc. 1977, Belknap and Kraft 1981, Kraft et al. 1983, Pearson et al. 1986, and Dix et al. 

2008). Using such circumstances, this study provides a methodology for managing conflicts between the 

preservation of prehistoric cultural resources and resource extraction by providing a basis for identifying 

and eventually mapping paleosurfaces with which significant cultural resources might be associated. The 

same methodology also provides a basis for mapping potential sand resources for use in coastal 

restoration.

Furthermore, from an academic, archaeological point of view, the need to reconstruct submerged 

paleolandscapes is multi-fold (Dix et al. 2008). First, for the accurate interpretation of the terrestrial 

record of the modern coastal plain, it is essential that archaeologists understand the contemporaneous 

paleogeography, including paleoshoreline locations and configurations, of now-submerged portions of 

prehistoric coastal plains. Second, the archaeological, paleontological and paleobotanical potential of the 

submerged paleolandscapes must be understood both in terms of primary cultural deposits that are 

preserved as a result of submergence and secondary cultural materials that have been re-worked as the 

result of the latest cycle of exposure and marine inundation. 
11

2 Methodology

2.1 Background and Archival research

The first phase of this study consisted of collecting the existing subsurface data for the Southwest 

Louisiana Outer Continental Shelf (OCS) study area (Figure 2.1). The data consisted of data collected by

Louisiana Geological Survey (LGS), Texas A&M University (TAMU), the Bureau of Ocean Energy 

Management (BOEM), Louisiana Department of Natural Resources (LaDNR), Louisiana Coastal 

Protection and Restoration Authority (CPRA), and US Geological Survey (USGS) since the early 1980s 

for this region. It included LGS and LaDNR sand resource survey studies; CPRA data for site specific 

sand borrow area prospecting; GIS files from Heinrich et al. (2015); USGS seismic data collected in 1994 

and 1995; 2010 TAMU-BOEM vibracore data from the Sabine incised valley area; Rice University data; 

block surveys conducted for geologic hazards and cultural resource surveys; and bathymetry, and offshore 

platform and onshore geotechnical borings. Key geologic data (sediment cores and geotechnical borings)

from these data were used to constrain two-way travel time compared to true depth relationships and in 

evaluating stratigraphic and other interpretations made from the seismic data. 

Figure 2.1. Southwest Louisiana Outer Continental Shelf Study Area.
12

2.2 Literature Review and Synthesis 

In the next phase of this study, an intensive examination was made of published literature concerning the 

geoarchaeology of the Louisiana OCS region and adjacent coastal plain. This material included peerreviewed literature, theses, dissertations, gray literature cultural resource management reports, and BOEM 

reports. Where found, useful paleogeographic and isopach maps were copied and later digitized for use in 

this project. The information gained from these sources was used to interpret the paleogeography, 

depositional-erosional history, and sedimentary processes associated with the accumulation of sediments 

within the study areas. This information, along with available core and seismic data, was also used in the 

construction of cross-sections and correlation of allostratigraphic units from onshore to offshore and to 

interpret depositional facies, potential for sand resources, and the nature of contacts between 

allostratigraphic units. Finally, conceptual models for paleosurface preservation potential and paleovalley 

sand resource occurrence on the northern Gulf of Mexico OCS were produced based on the available data 

and concepts of sequence stratigraphy.

BOEM requires that surveys in the Gulf of Mexico use US measurement units (feet and inches). Because 

of the prevalence of oil and gas industry-related research in the Gulf, many of the references, maps, and 

data were initially obtained and/or reported in feet. Formatting guidelines for BOEM studies require 

International or metric units as the primary convention. All measurements in this report, therefore, are 

given in metric (SI) units. In most cases, the original unit was feet; metric conversions were calculated 

specifically for this report.

2.3 Offshore Lease Block Hazard Survey Reports Database Development

An integral part of this project was development of a GIS geodatabase for visual representation, data 

inventory tracking, and computational support. The GIS application used in this project is Environmental 

Systems Research Institutes (Esri®) ArcMap® v10.1. Digital map data were loaded into GIS for 

interpretive analysis. Base and thematic data were acquired from state and federal sources. See Appendix 

A for a list of layers. The majority of the GIS work was to convert portable document Format (PDF) 

raster maps to digital spatial data for incorporation into the GIS. These documents were included in Lease 

Block Hazards Survey Reports submitted to BOEM for compliance. The reports contain maps produced 

from data interpretations and no original data products (e.g., navigation files, geophysical profiles). Three 

hundred and fifty interpreted hazards and cultural resource maps were made available for this study area 

and spanned over three decades of survey data, dating to the late 1970s. Accompanying 28 of the PDF 

files were original computer aided design (CAD) files used to generate the PDF maps for inclusion in the 

GIS project. 

The process of developing GIS data from Portable Document Format (PDF) Geophysical Hazards Maps 

was challenging because of inconsistent reporting formats, drafting styles, and nomenclature used to 

describe geologic features, and compilation of the data is detail-intensive. Initial steps involved

determining which PDF map files were relevant to this study on shelf paleovalleys and paleochannels The 

relevant information consists of numerous paleovalleys and a few scattered paleochannels and channel 

belts that were mapped as a part of individual lease block hazard surveys (Fig 2.1). These surveys often 

only cover a 3 x 3 NM area of the shelf coincident with an OCS lease block that must be cleared for 

hazards and potential cultural resources prior to development by the lessee. Relevant PDF map files were 

exported from Adobe® Acrobat® as Tagged Image File Format (TIFF) files. A spreadsheet was created 

to document the content of each PDF file and track whether it was digitized for the study. 

Each TIFF file was assigned spatial reference defining where the features are located in the real world. 

This task was accomplished by using the georeferencing toolbar in ArcMap®. One hundred and twentyone TIFF files were digitized within the study area in East Cameron, West Cameron, and Sabine Addition 

Areas offshore western Louisiana and Texas. After georeferencing each map, line or point, shapefiles 
13

were created with ArcCatalog applying the respective spatial reference. Digitizing the lines and/or points 

was done manually while in edit mode on each file. Much effort was devoted to ensuring features were 

represented as they appear on the maps. The vectorization tool in ArcMap®, which automatically traces 

linework, was of little use because of extraneous line data on the maps such as track lines or deeper 

geologic features not of interest to this study. Digitized lines were further developed into polygons using 

the Polyline to Polygon tool in ArcToolbox®. These polygon data typically represent the paleochannel

and/or paleovalley edges, while the point data represent the tops and bases of the channels. 

Attribute tables were developed after digitization of the features. Unfortunately, most of the maps were 

unique and shared little legend continuity, which created issues with nomenclature development. This is 

likely a result of the different interpreters and interpretation dates and lack of an accepted standard for 

geologic feature nomenclature. The descriptions are recorded in the attribute tables (Table 1) along with 

other pertinent data. These include the lease block number assigned by BOEM. Descript column describes 

the attribute as defined in the maps legend. As mentioned, there are many different descriptions within 

these geophysical maps. The tables are populated with the descriptions as listed in Appendix B. 

In total, 350 offshore-block hazard surveys were reviewed of which 122 were digitally synthesized and 

228 were found to be unusable. Despite the problematic variability of the source information, its 

compilation and availability in one GIS should have value as a reference resource (Figure 2.2).

Figure 2.2. Data Distribution Southwest Louisiana Outer Continental Shelf Study Area.
14

Most of the mapped features in the final compilation appear to be smaller valleys of upland tributaries 

draining interfluves, though some portions of larger and/or wider features were identified: 1) in the 

western portion of the southwestern study area, correlating with parts of the Sabine and Calcasieu 

paleovalleys, and 2) in the eastern portion of the southwestern study area, for which we are aware of no 

obvious correlation with known features and the basis of which remains unclear (Figure 2.3).

Figure 2.3. Distribution of paleolandforms in southwest Louisiana Outer Continental Shelf Study 

Area, as mapped by OSC Hazard studies and Pearson et al. (1986). 
15

3. Northern Gulf of Mexico Geology and Sea Level History

3.1 Geologic Setting

The study area encompasses the northern and middle portions of the Southwest Louisiana Continental 

Shelf adjacent to the Southwest Louisiana shoreline from Joseph Harbor Bayou (Humble Canal), 

Cameron Parish west to about 6 km west of Sabine Pass along the shore of Chambers County, Texas. 

This study area lies within the coastal plain and continental shelf that forms the northwestern margin of 

the Gulf of Mexico. 

The Gulf of Mexico Basin, as used in this report, is a roughly circular structural basin that is 

approximately 1,500 km in diameter. It is filled with up to 15.1 km (9.4 mi) thick sedimentary deposits

ranging from Jurassic to Holocene in age. Locally, Triassic rift basins along its edges contain an 

additional 1 km or more of Triassic sedimentary units (Bryant et al. 1991, Salvador 1991). The Gulf of 

Mexico Basin is the result of rifting within the North American plate as it drifted away from the African 

and South American plates (Salvador 1991). As currently understood, the first phase consisted of Late 

Triassic–Early Jurassic (235–174 Ma) rifting that resulted in the opening of the Gulf of Mexico in a NWSE extension direction. This rifting also created a broad zone of thinned, continental crust along the 

northern margin of the Gulf of Mexico that now underlies the northern salt basins of Texas, Louisiana,

and Mississippi. However, it failed to produce in this area a parallel and contiguous zone of oceanic crust. 

During the Middle Jurassic, marine flooding of the proto-Gulf of Mexico basin deposited extensive 

evaporite, mainly salt, strata. After the accumulation of thick evaporites, a second phase of Late Jurassic 

(156–145 Ma) extension occurred along a highly arcuate, slow spreading ridge system. This ridge system 

created a large expanse of salt-free, oceanic crust that underlies the deepwater shared by the US, Mexico,

and Cuba (Nguyen and Mann 2016). During the Early Cretaceous, carbonate platforms and shelf margins 

rimmed the Gulf of Mexico and thick accumulations occurred along its northwestern rim (Winker and 

Buffler 1988). Carbonate sedimentation ceased as the result of continental-scale drainage reorganization 

that occurred between Middle Cretaceous, Aptian-Cenomanian, and Paleocene. This reorganization 

expanded the area drained into the Gulf of Mexico from largely the Appalachian Mountains to include 

fluvial systems of western and midcontinental North America that were routed either directly into the 

Gulf of Mexico or the Mississippi Embayment (Blum and Pecha 2014). As a result, km-thick sequences 

of sandstones and mudstones buried the Cretaceous carbonates and extended basinward to the base of the 

modern continental slope along the northwestern rim of the Gulf of Mexico over the Cenozoic (Bentley et 

al. 2016).

The current, Quaternary coastal plain of Texas and Louisiana consists of the strata that were deposited by 

fluvial systems of various types, sizes, and sediment composition (Young et al. 2012). Despite these 

variations, these fluvial systems can be divided into extrabasinal and intrabasinal rivers. Extrabasinal 

rivers are rivers that have large drainage basins that extend well beyond the coastal plain. Over the 

Cenozoic, such rivers have persistently occupied major embayments within the coastal plain as they 

currently do. For example, the Rio Grande, Colorado-Brazos, and Mississippi rivers, respectively, occupy 

the Rio Grande, Houston, and Mississippi embayments. The point where an extrabasinal river enters the 

coastal plain is stable owing to the entrenchment of its valley across the slightly uplifted margin of the 

coastal zone (Winker 1979). Gulfward from this entry point, extrabasinal fluvial systems freely migrate 

laterally across the coastal plain and, thus, construct alluvial plains composed of sand-rich channel-fill 

facies and mud-rich floodplain facies (Figure 3.1) (Morton and McGowen 1980, Galloway 1981; Blum 

and Price 1998, Anderson and Fillon 2004).
16

Figure 3.1. Diagrammatic interpretation of Quaternary depositional systems of the Texas Coastal 

Plain. 

Reprinted with permission from Young et al. (2012).

In contrast, intrabasinal rivers are rivers that have drainage basins marginal to and within the coastal 

plain. Within both East Texas and Southwest Louisiana, the overall drainage pattern and individual 

courses of intrabasinal rivers are strongly controlled by the relict antecedent topography of major 

extrabasinal rivers and coast-parallel terraces. In the Houston region, intrabasinal rivers and streams flow 

roughly parallel to highstand, interglacial meander-belt ridges within interridge lowlands and along the 

gulfward edges of coastal plain segments (Van Siclen 1985). In Southwest Louisiana, the courses of the 

Mermentau River and its tributaries occupied interchannel lows that lie between relict Red River fluvial 

and beheaded deltaic distributary ridges. 

The Quaternary alluvial aprons grade basinward into deltaic and shore-zone depositional systems. In the 

Rio Grande and Houston embayments, the major extrabasinal rivers have constructed sand-rich deltaic 

headlands on the shore of the modern coastal plain. Within and gulfward of the Mississippi embayment, 

the Mississippi River has constructed a major deltaic plain. Intrabasinal rivers empty into underfilled 

estuaries with bayhead deltas at their apex. Elsewhere along the Gulf of Mexico, coastal plain, lagoon and 

barrier island systems, and the Louisiana Chenier Plain form the coastline (Morton and McGowen 1980,

Galloway 1981).

The Cenozoic strata underlying the northwest coastal plain of the Gulf of Mexico consist of a 

heterogeneous mixture of sand, sandstone, mud, and mudstone that lacks any distinctive lithologic 

subdivisions (Galloway et al. 1991). These sediments have been subdivided using a combination of: 1) 

biostratigraphic zonation; 2) fluvial-deltaic depositional models based on Quaternary examples; and 3) the 

definition of regionally cyclic depositional episodes (Galloway et al. 2000, Young et al. 2012). The 

extinction points of biostratigraphic zonation are known as “last occurrence or datum top” of calcareous 

nannofossils and planktonic foraminifers and other marine microfossils (Galloway et al. 1991, Lawless et 

al. 1997, Fillon and Lawless 2000). Because of the absence of marine fossils in fluvial sediments, 

stratigraphic subdivisions are extended downdip from outcrop using volcanic ash layers, well log 

correlation techniques, and limited nonmarine (vertebrate faunal) biostratigraphy (Lundelius 1972,

Tedford and Hunter 1984, Baskin and Hulbert 2008, Young et al. 2012). 
17

3.2 Sea Level Change: An Overview

Change in relative sea level is the result of the complex and independent interaction between the 

movements of two different surfaces: the surface of the sea and either the land surface or sea floor. The 

surface of the ocean is spatially and temporally variable related to total mass of ocean water compared to

water that is sequestered as continental ice sheets, the eustatic component. Tectonic deformation of either 

land or the sea floor is spatially and temporally variable and arises, in part, due to changes in ice loading 

of the crust such as during a during a glacial cycle, subsidence induced by sediment compaction, or 

tectonic loading (Douglas 2001, Dix et al. 2008).

As discussed by Dix et al. (2008), the diverse factors that control fluctuations in both the elevation of the 

either the land surface or sea floor and either the sea or ocean surface interact on a wide range of spatial 

and temporal scales. They vary from astronomical factors that operate from outside the Earth, such as 

variations in the planet’s angular velocity, to other factors that operate on the Earth’s surface, like the 

global distributions of glaciers and meltwater. Finally, there are factors working within the Earth’s 

interior, such as displacements of mantle material (Pirazzoli 1993, Douglas 2001). These factors are often 

interconnected and they operate in conjunction with each other. The dominant influence on relative sea 

level change at a specific place and time depends on the temporal and spatial scale on which it is observed 

(Dix et al. 2008).

3.2.1 Eustatic controls on sea level

During the Quaternary, the dominant mechanism responsible for sea-level changes has been the cyclic 

addition and subtraction of water from global oceans from the progressive buildup and decay of 

continental-scale ice sheets in response to Milankovitch forcing insolation changes. Minor input and 

outputs to the volume of water stored in the world’s oceans and seas have come from atmospheric water, 

rivers, lakes, peat bogs, and so forth as part of the hydrological cycle. However, variations in storage in 

these sources can contribute only relatively little–a layer of water some 58.3 cm thick evenly distributed 

worldwide–to present ocean volume (Dix et al. 2008). 

Throughout the Quaternary, the interplay between variations in the Earth’s orbit and axial tilt; ice sheet 

dynamics; and ocean circulation have resulted in a long-term cycle of climate change consisting of 

alternating glacial and interglacial phases. These cycles were about 41,000 years in the early Pleistocene 

and Pliocene and about 100,000 years from about 800 ka BP to present. The removal of waters from 

global oceans during the formation of continental ice sheets resulted in sea level lowstands and melting of 

continental ice sheets resulted in sea level highstands (Zachos et al. 2001, Lambeck et al. 2002, Dix et al. 

2008).

Such relative changes in global sea level, which were on the order of 100 m, were termed “eustatic” by 

Suess (1906). Thus, eustatic sea-level change is defined as the global shift in the height of the ocean 

surface that occurs, in the absence of gravitational and rotational effects, ocean dynamics, and solid Earth 

deformation, due to the input of meltwater to the ocean. Because eustatic sea-level change is directly 

equivalent to the changing mass of the global ice sheets, it provides an independent method for calibrating 

the marine oxygen isotope record and the total volume of continental ice at the Last Glacial Maximum 

(Farrell and Clark 1976, Dix et al. 2008, Whitehouse and Bradley 2013). Since the end of the Last Glacial 

Maximum at about 20 ka BP, sea-level change has been dominated by the addition of meltwater into the 

oceans associated with the retreat and shrinkage of the major ice sheets. Thus, it has been postulated that

locally, relative sea-level change in tectonically stable regions far from the major ice sheets, known as farfield sites, approximates eustatic sea-level change since the Last Glacial Maximum (Milliman and Emery

1968).
18

3.2.2 Vertical displacement of oceanic and continental crust

Active displacement, both uplift and subsidence, of the Earth’s crust is a quite common occurrence 

anywhere on the Earth’s surface, including the Gulf of Mexico coastal plain and continental shelf. Both 

subsidence and uplift, respectively, are associated with isostatic adjustments as the result of increased or 

decreased loads on the earth's surface. It is called “glacio-isostasy” when the load is ice, “volcanoisostasy” when it is extruded magma, “sedimento-isostasy” when it is sediment deposition or erosion, and 

“hydro-isostasy” when it is a layer of water (Pirazzoli 1991). Finally, either temperature and/or density 

changes of materials inside the Earth’s interior can cause uplift and subsidence of the Earth’s crust. It is 

called thermo-isostasy when temperature is the predominant factor, e.g., near a hot spot (Pirazzoli and 

Grant 1987). Along the northern Gulf of Mexico, for the late Quaternary, glacio-isostasy and sedimentisostasy have been demonstrated to influence relative sea-level (Wolstencroft et al. 2014). 

3.2.1.1 Glacio-isostasy 

The cyclic transfer of large volumes of water from the oceans to land ice and back that characterized the 

Quaternary changes not only created the eustatic component of sea level, it also resulted in increased or 

decreased loads on the Earth’s surface as continental ice sheets grew and shrank in area and volume. This 

redistribution of surface water mass is large enough to have caused a significant, cyclic glacio-isostatic 

deformation, known as “glacial isostatic adjustment” of the solid Earth. It is one of the key processes that 

contributed to glaciation-induced sea level change. During the glacial maximums, ice sheets reached 

thicknesses of several kilometers in places. The accumulated mass of glacial ice sheet was sufficient to 

depress the underlying crust by the order of hundreds of meters. Peripheral to areas of depression that 

underlie the ice sheets, there was corresponding uplift of the crust to form a forebulge in front of it 

(Chappell 1974, Peltier and Andrews 1976, Mitrovica and Peltier 1991, Milne and Shennan 2013).

During the retreat and disappearance of the ice sheets from North America and Eurasia during the early to 

mid-Holocene, there was, and still remains, a significant uplift in these regions. This is the result of the 

relatively high viscosity of mantle material and the resulting slow rate at which the Earth’s crust is 

attaining isostatic equilibrium following the removal of ice sheets. There is also a corresponding belt of 

relatively slow isostatic subsidence that result from the collapse of the forebulges surrounding the regions 

depressed by glacial ice sheets (Peltier and Andrews 1976, Mitrovica and Peltier 1991, Lambeck 1993,

Milne and Shennan 2013).

It has often been assumed that glacial isostatic adjustment was insignificant within the Southwest 

Louisiana-Southeast Texas coastal plain and continental shelf (Gagliano et al. 1982; Anderson and Bissett 

2015). However, modeling of glacial isostatic adjustment demonstrates that it had some effect on relative 

sea level (Simms et al. 2007, Milne and Shennan 2013). The magnitude of the effect of glacial isostatic 

adjustment on relative sea level, as discussed later, can be seen in the maximum MIS 2 lowstand of 

relative sea level in the northern Gulf of Mexico of between -80 and -90 m when the eustatic sea level 

was about -120 m from between about 29 and 18 ka (Waelbroeck et al. 2002).

3.2.1.2 Hydro-isostasy and continental levering

Another process that affects relative sea level by causing significant deformation of the Earth’s crust is 

hydro-isostasy. It is the deformation of the Earth’s crust as a coastal plain is submerged beneath rising sea 

level. The weight of the water loading the crust causes it to subside even further. In addition, as the crust 

is loaded by rising sea level, it not only subsides, it also forces mantle landward beneath crust, causing it 

to be uplifted. This uplift is called “continental levering.” The effects of hydro-isostasy and associated 

continental levering is most discernable in regions, “far-field” regions, far removed from the major ice 

centers where the effects of glacial isostatic adjustment are considerably reduced, as shown by the small 

uplift around the coasts of Australia and Africa (Clark et al. 1978, Milne and Shennan 2013).
19

3.2.1.3 Ocean syphoning

During deglaciation, both hydro-isostasy and glacial isostatic adjustment increase the volume of the 

world’s oceans by lowering sea floor and deepening the oceans within collapsing forebulges and 

subsiding continental shelves. This deepening of the ocean floor will increase volume of the ocean and its 

capacity to store water in these areas and draw (“syphon”) sea water away from more stable far field, 

largely equatorial, areas and cause a fall in relative sea level within them. The magnitude of the global 

mean sea-level fall due to syphoning during deglaciation can be a few 10s of meters from the Last Glacial 

Maximum to present (Mitrovica and Peltier 1991, Mitrovica and Milne 2002).

3.2.1.4 Sedimento-isostasy 

Another source of tectonic displacement of the crust within the northwestern Gulf of Mexico is 

sedimento-isostasy. Before anthropogenic modification o f the Mississippi River, its tributaries and 

drainage basin steadily supplied large sediment loads to the continental margin with little long-term 

sediment storage within the lower Mississippi River Valley. This process has resulted in the construction 

and extension of the Mississippi River Delta Plain southward into the Gulf of Mexico Basin, which began 

in the early Paleocene. The loading of the continental shelf with fluvial sediments was greatly accelerated 

during Pleistocene glacial periods, most recently during MIS 4 through MIS 2. Later, during the 

slowdown in deglacial sea level rise about 8000 ka, near-shore aggradation of this fluvial sediment load 

allowed the development of the Mississippi River Deltaic Plain (Coleman and Smith 1964). The weight 

of this load of accumulating deltaic sediments induced subsidence due to downward flexure within the 

underlying crust (Bowie 1927, Fisk and McFarlan 1955, Blum et al. 2008). In this process, the Earth’s 

crust deforms elastically in that it deforms under the stress, but it reverts to its original form after the 

strain ceases. In contrast, the mantle acts as a highly viscous fluid and permanently deforms under the 

stress of the accumulated sediments. 

The accumulation sediments on the underlying lithosphere (i.e., the crust and upper mantle) causes mean 

downward surface displacement of the underlying lithosphere, i.e., the crust and upper mantle. If the 

mean increase in elevation due to sediment aggradation over an affected area is less than the downward 

displacement of the surface then subsidence and eventual submergence will occur. If the mean 

accumulation of sediment over an affected area is equal to or greater than the downward displacement of 

the surface the surface of the affected area will remain stable and the excess sediment will be transported 

seaward and eventually result in gulfward progradation and forced regression of the shoreline. Sediment 

loading differs from sediment compaction because it refers to subsidence in the material underlying a 

sediment load rather than subsidence resulting from volume decrease within the sediments themselves 

(Fisk and McFarlan 1955, Coleman and Smith 1964, Diegel et al. 1995, Dokka et al. 2006).

The process of growth faulting has also been attributed, in part, to sedimento-isostasy resulting from the 

construction and extension of the Mississippi River Delta Plain southward into the Gulf of Mexico Basin. 

This process began loading with fluvial sediment in the early Paleocene. It has been argued that flexure of 

the continental crust has reactivated growth faults with general east-west orientations, perpendicular to the 

direction of delta growth. Also, growth faults are created as the delta progrades down an inclined surface. 

In time, the prograding delta front overextends and detachment and the formation of fault zones, which 

slip by breakaway, and gravitational slumping, occurs (Berman 2005, Diegel et al. 1995, Dokka et al. 

2006).
20

Although the Mississippi River Delta is well recognized for its associated sedimento-isostasy subsidence, 

there is evidence of sedimento-isostasy subsidence association with the valley fills and deltas of the 

Colorado, Brazos, and Rio Grande rivers. As noted by Graf (1966) and Winker (1990), the elevation of 

beach ridges associated with the Ingleside shoreline systematically varies and can be interpreted as 

reflecting subsidence associated with the valleys of Brazos and Colorado rivers and the Mississippi 

alluvial valley and delta and uplift between them.

3.2.3 Halokinesis and surface deformation

Halokinesis, salt tectonics, also can cause deformation, either subsidence or uplift of the surface of the 

Gulf of Mexico coastal plain and sea bottom. Large pillars of salt intrude upward through kilometers of 

the largely unlithified sediments that underlie the surface of the coastal plain and sea bottom as the result 

of buoyancy effects. The upward relative movement of salt domes can result in the uplift of local, roughly 

circular areas, e.g., the Five Islands and the subsidence of a moat-like area around an uplifted area. The 

movement of salt domes can create radial faults and downthrown grabens. Subsidence resulting from 

halokinesis would occur very slowly over geologic timescales (> 1000 yrs). The spatial scale would either 

be the areal diameter of the diapir, 1 to 100 km2

, or regional as halokinesis causes large blocks to slide 

gulfward along regional growth faults (Berman 2005, Diegel et al. 1995, Dokka et al. 2006, Reed and 

Yuill 2009).

3.2.4 Sediment compaction and subsidence

In coastal environments, newly deposited fluvial and marine sediment may compact in time as the result 

of various physical, biological, and chemical processes (van Asselen et al. 2009). Within coastal 

Louisiana, including the Mississippi River Delta, physical compaction is the most commonly cited 

compaction-related cause of subsidence. The process of compaction includes both compression and 

consolidation of the sediments. Compression is the decrease in soil volume due to a restructuring of 

internal grain alignment as a result of an applied stress. This restructuring of the sediment causes the 

sediment to become more tightly packed. Unlike consolidation, compression occurs in relatively short 

timescales upon application of an applied load and is generally controlled by the geotechnical properties 

of the soil. Consolidation relates to the time-dependent expulsion of pore water in response to an applied 

stress, which reduces the internal pore pressure causing pore collapse. Unlike compression, consolidation 

is a gradual process that is controlled by soil-water interactions. In sediments composed of peat and other 

organic-rich sediments, biological (microbial decay of organic material) or chemical (oxidation of organic 

carbon) processes can result in significant compaction (van Asselen et al. 2009, Reed and Yuill 2009).

3.3 Relative Sea Level History 

Beginning about 2.6 million years ago, continental ice sheets became a prominent feature in the Northern 

Hemisphere. Later they became widespread elsewhere over the Earth’s surface in response to the cooling 

global climate. During the Quaternary Period, which is divided into the Pleistocene and the Holocene 

epochs, these continental ice sheets episodically covered millions of square kilometers and many were up 

to 3 km thick. The water that produced the huge quantities of snow that created these continental ice 

sheets came from the world’s oceans. The result was a global drop in sea level during glacial epochs that 

lowered global (eustatic) sea level by up to 130 m (Figure 1.1). The periodic drop in global sea level 

caused repeated exposure of part of the coastal plain of what is now the Gulf of Mexico continental shelf. 

The carbonate platforms of Florida and Yucatan were also periodically exposed with the development of 

karstic surfaces with innumerable sinkholes. Of these glacial–interglacial cycles, this study is concerned 

only with the last one, from MIS 5 to MIS 1, because that is when landforms and landscapes that form the 

background to this study were created (Figure 1.2) (Bentley et al. 2016).
21

3.3.1 MIS 5, Last Interglacial

The last interglacial period, MIS 5, spans the period from about 130 to 71 ka. This period was the last 

interval of the Pleistocene when global ice volumes were small, and ice-equivalent sea levels oscillated 

from +6 to -45 m, with a mean value of -27 m. Regionally, the MIS 5 sea levels departed from these 

globally averaged values because of GIA contributions and vertical tectonic displacements (Waelbroeck 

et al. 2002). During MIS 5e, ice-equivalent sea levels were 6–9 m higher than present (Kopp et al. 2009

and 2013, Dutton and Lambeck 2012, Anderson et al. 2014, Bentley et al. 2016).

Offshore shorelines for periods of sea level fall after MIS 5e are poorly constrained. Simms et al. (2009) 

concluded that the MIS 5a paleoshoreline occurs at -11 ± 2 m water depth on the east Texas continental 

shelf. They also refined earlier estimates by Rodriguez et al. (2000) of the depth of the MIS 3 

paleoshoreline and concluded that it might have been as high as -20 m, after correcting for local 

glaciohydro-isostatic effects (Figure 1.2).

3.3.2 Last Interglacial highstands, MIS5, Gulf of Mexico

During the early part of the last interglacial, MIS 5e, sea level was at its maximum highstand position, a 

few meters above present sea level. It was at this time that an ancient beach-ridge complex, referred to as 

the Ingleside Paleoshoreline, was created. It marks the location of the MIS 5e shoreline from eastern 

Louisiana to just south of the city of Corpus Christi, Texas. In East Texas and Louisiana, the Ingleside 

Paleoshoreline is located 35 to 60 km inland of the modern shoreline. Its preservation demonstrates that it 

was never subjected to either transgressive or regressive shoreface erosion and it likely marks the 

landward extent (maximum flooding shoreline) of the Gulf of Mexico during MIS 5 (Graf 1966;

Barrilleaux 1986, Winker 1990).

The Western Louisiana segment of the Ingleside Paleoshoreline, the Houston Ridge, is a 32 km-long, 

narrow, 500 to 900 m wide, east-west trending ridge. The crest of this ridge ranges in elevation between 

7.5 and 11 m above mean sea level and lies 2 to 4.5 m above the surface of the adjacent Beaumont 

Alloformation. The development of innumerable pimple mounds has extensively modified the surfaces of 

both the Houston Ridge and other parts of the Beaumont Alloformation. However, using LIDAR DEMs, 

remnants of beach ridges can be observed within the eastern end of the Houston Ridge. The Ingleside 

Paleoshoreline is absent locally where removed by fluvial processes (Figure 3.2) (Heinrich 2007).
22

Figure 3.2. Geomorphic map of Southwest Louisiana Pleistocene paleochannels. 

Compiled and generalized from McCulloh and Heinrich (2002), Saucier and Snead (1989), and various

published and unpublished 1:100,000 scale Louisiana geologic quadrangles (Heinrich 2005a).

Despite initial attempts to directly date this paleoshoreline providing mixed results (e.g. Shideler 1986,

Otvos and Howat 1997, Blum et al. 2002), the Ingleside Paleoshoreline has long been regarded as

representing the ancient shoreline of the last interglacial highstand (Price 1933, Graf 1966, Barrilleaux 

1986, Morton and Price 1987, Paine 1993, Anderson et al. 2004). More recent research by Simms et al. 

(2013) yielded six optically stimulated luminescence ages for segments of the Ingleside Paleoshoreline 

near Smith Point east of Galveston Bay and Port O'Conner, Texas. The ages obtained by Simms et al. 

(2013) cluster into groups of 119 ka–127 ka BP, 54 ka–57 ka BP, and 1.3 ka BP. The two groups of 

younger ages appear to be related to two periods of regional eolian reworking of coastal plain surface 

(Otvos 2004, Forman et al. 2009, Simms et al. 2013). The ages of 119–127 ka BP support earlier 

interpretations that the Ingleside Paleoshoreline is an MIS 5e shoreline deposit (Figure 1.2) (Anderson et 

al. 2004 and 2016).

After the last interglacial MIS 5e, global sea level experienced significant oscillations. During MIS 5, 

these fluctuations were MIS 5d to MIS 5a, of which MIS 5c and MIS 5a correspond with each interstadial 

sea level highstands. From the interglacial highstand of MIS 5e, sea level fell abruptly about 115 ka BP 

down to -60 m. Afterwards, sea level rose to about -15 m around 103 to 98 ka BP and then fell again to 

about -60 m during MIS 5b. During MIS 5a about 85 ka BP, sea level rose again 15–25 m of its present 
23

level. Sea level was about 65 m below present at the end of MIS-5a about 75 ka BP (Lambeck and 

Chappell 2001, Lambeck et al. 2002, Cutler et al. 2003, Potter et al. 2004, Hopley et al. 2007).

The surface characteristics of the Beaumont Alloformation north of the Ingleside Paleoshoreline varies 

greatly. Within Southeast Texas, its surface consists of a mosaic of imbricated offlapping alluvial and 

deltaic plains that exhibit relict fluvial ridges and paleochannels of meander belts and deltaic 

distributaries. Since their abandonment, an inversion of topography has occurred in which modern 

drainages have come to occupy the interchannel areas between fluvial ridges and deltaic distributaries and 

the boundaries of individual offlapping coastal plains. Based on stratigraphic and geomorphic 

considerations, except for terraces along modern drainages, these features are inferred to predate MIS 5e 

(Van Siclen 1985 and 1991, Aronow 1971). Within Southwest Louisiana, the narrow belt of coast-parallel 

terrace lying between Ingleside Paleoshoreline and where it onlaps the surface of the Lissie Alloformation 

consists of a typically flat plain that is featureless except for the occasional paleochannels of a small-scale 

meandering channel belt (Heinrich 2007).

South and east of the Ingleside Paleoshoreline within the Louisiana Coastal Plain, the surface of the 

Beaumont Alloformation is characterized by two different groups of relict fluvial-deltaic depositional 

topography (Figure 3.2) (Fisk 1948, Saucier 1977, Aronow 2004). East of Bayou Choupique, the coastparallel terrace of the Beaumont Alloformation exhibits paleochannels and depositional ridges of Red 

River channel belts and deltaic distributaries. These features range in degree of preservation from the 

well-defined meander loops to either fragments of channels and detached, discontinuous loops or 

featureless distributary or fluvial ridges. These relict fluvial-deltaic features, especially east of the 

Calcasieu River, impart a northeast-southwest grain to the drainage patterns and, similar to the area in 

Texas discussed above, there has been an inversion of topography with the modern drainage now 

occupying interchannel lows between paleochannels. Some of the paleochannels have distinctly incised 

their channels, often leaving beheaded deltaic distributaries on both sides of them. There is an apparent 

lack of continuity in trend and scale between the paleochannels and channel belts observed on the surface 

of the Beaumont Alloformation and the “fluvial channels” of Suter (1986 and 1987); the “distributary 

channels” (“incised valleys”) of Wellner (2001); and the “channel belts” (“channel complexes”) of 

Anderson et al. (2016). Finally, on the surface of the Beaumont Alloformation, a lower relief and more 

poorly defined set of possible beach ridges extend for a distance of about 15 km from the east end of 

Houston Ridge at Sam Houston State Park to the west valley wall of the Calcasieu River.

The western portion of the terrace surface of the Beaumont Alloformation exhibits innumerable poorly 

preserved, rather obscure, and fragmentary meander loops of paleochannels. These fragmentary 

paleochannels become better defined and more continuous to the northwest, towards Starks, Louisiana. 

They also impart a northwest-southeast grain, which controls the minor drainages, to the landscape. 

Bernard and LeBlanc (1965) and Aronow (2004) infer that this portion of the terrace surface of the 

Beaumont Alloformation was laid down by the Sabine River. 

The terrace surface and surficial sediments south and east of the Ingleside Paleoshoreline postdate the 

MIS 5e highstand based on cross-cutting relationships. In addition, the contemporaneous incision of 

fluvial channels, dissection of deltas and beheading and stranding of their distributaries are indicative of 

forced regression brought on by falling relative sea level. 

As summarized by Shen et al. (2012), the age and correlation of the coast-parallel surface and associated 

sediments that form the upper contact of the Beaumont Alloformation have been the subject of 

considerable speculation. The first optically stimulated luminescence ages for the uppermost, coastparallel sediments of the Beaumont Alloformation by Otvos (2005) yielded ages of MIS 5 to MIS 3. 

However, as discussed by Shen et al. (2012), Otvos (2005) lacks the necessary documentation regarding 

either the dating methodology or the stratigraphic context of dated samples. Commenting on the 
24

inconsistencies in these dates noted by Heinrich (2006), Shen et al. (2012) proposed that many of these 

optically stimulated luminescence ages underestimate the age of the overbank deposits due to pedogenic 

and biogenic disturbance. Well-documented by Miller et al. (1985), Schumacher et al. (1988), and Autin 

and Aslan (2001), the surficial fluvial deposits of the Beaumont Alloformation are capped by a welldeveloped paleosol and are mixed with the overlying Peoria Loess in the eastern part of Southwest 

Louisiana. In this part of Southwest Louisiana, this surface was subaerially exposed for at least thousands 

of years during which significant pedoturbation of the uppermost few meters of these sediments occurred 

(Miller et al. 1985, Autin and Aslan 2001). In the western part that lacks a loess cover, the surface would 

have been exposed to pedoturbation for tens of thousands of years and at least to two periods of regional 

eolian reworking of coastal plain surface (Otvos 2004, Forman et al. 2009) as evidenced by optically 

stimulated luminescence ages published by Simms et al. (2013) from the Ingleside Paleoshoreline. In 

addition, there likely was significant influx to and mixing of wind-blown dust into the surface of the 

Beaumont Alloformation. As a result, there certainly has been significant pedoturbation and addition of 

eolian dust to the uppermost few meters of its sediments since the surface of the Beaumont Alloformation 

was abandoned.

According to Shen et al. (2012 and 2013), the upper Beaumont Alloformation aggraded in association 

with a MIS 5a sea-level highstand and subsequent abandonment of the coastal plain is not only 

demonstrated by their optically stimulated luminescence ages, but also the optically stimulated 

luminescence ages from and stratigraphic position of the younger Lafayette Meander Belt of the 

Avoyelles alloformation (Mateo 2005). Also, as noted by Shen et al. (2012 and 2013), the longitudinal 

profiles of the Mississippi River valley train from MIS 4 and 3 occur at elevations about 10 m below the 

MIS 5a long profile of the surface of the Beaumont Alloformation (Rittenour et al. 2007). Finally, the 

longitudinal profiles of fluvial terraces of the MIS 4 and 3 Deweyville Allogroup along the Calcasieu, 

Mermentau, and Sabine rivers also lie well below the elevation of the surface of the Beaumont 

Alloformation (Blum et al. 1995 and 2013, Heinrich et al. 2002 and 2003, Heinrich 2005b and 2006). 

Optically stimulated luminescence ages from the Freeport Rocks Bathymetric High, off of the Central 

Texas coast, show that their age is about 91 ka BP. Based on their relationship to former relative sea 

levels, Simms et al. (2009) concluded that MIS 5a sea levels within the northwestern Gulf of Mexico 

were around −11±2 m at this time. This is consistent with an estimated MIS 5a sea level highstand of 

around -7 to -9 m about 98 ka BP for the southwestern Louisiana continental shelf according to the 

ICE5G_120550_Full model of Martini (2015).

3.3.3 Falling-stage sea level: MIS 4 and 3

After MIS 5a, global sea level fell irregularly through MIS 4 and 3 into the Last Glacial Maximum, MIS 

2. During this period, global sea level rapidly fluctuated, with cycles of ~15 m amplitude recurring every 

6–7 ka. Global sea level rose to close to -40 m midway through MIS 4, then fell to about -70 m by 65 ka 

BP, and rose again to about - 50 m by the start of MIS 3. During MIS 3, 70–30 ka BP, sea level fluctuated 

between highstand and lowstands between -45 m and -90 m several times (Figs 1.1 and 1.3; Chappell, 

2002, Lambeck et al. 2002, Anderson et al. 2016).

Within the southwest Louisiana continental shelf, the ICE5G_120550_Full model of Martini (2015) 

estimates that during MIS 4, relative sea level fell to just less than -50 m about 56 ka BP from around -7 

m to -9 m about 98 ka BP (MIS 5). For the MIS 4 lowstand, the same model estimates that the highest 

that relative sea level rose during the MIS 3 highstand was about -45 m and the lowest that it dropped was 

about -90 m. 
25

Based on stratigraphic studies, the older (“Early Wisconsin”) channel belts and associated shelf-margin 

deltas (mapped by Suter 1987) formed during an informal “falling-stage sea level” period that consists of 

MIS 5d to 5a, 4 and 3 (Figures 3.4 and 3.5; Coleman and Roberts 1988a and 1988b), Wellner et al. 2004,

Anderson et al. 2004, 2014, and 2016). Suter and Berryhill (1985), Berryhill (1987), and Suter (1987) 

conducted the original, detailed studies of these “falling stage” channel belts and shelf margin deltas

mapping fluvial channel belts on the western Louisiana and Texas continental shelves and concluded that 

they formed during “Early Wisconsin” time. 

Figure 3.3 MIS 5–3 channel belts of the western Louisiana continental shelf, as modified from 

Suter and Berryhill (1985) by Anderson et al. (2016).

Anderson et al. (2015) subdivides them into three channel belts complexes. The eastern channel belt 

complex merges seaward with the Mississippi shelf margin delta. A second, narrower western channel 

belt complex merges with MIS 3 Western Louisiana Delta (WLD). The western most channel belt complex 

they assign to the the Calcasieu and Sabine paleovalleys. MRCC = Mississippi River channel complex; 

WCC = Western channel complex. Reprinted from Anderson et al. (2016).
26

Figure 3.4. Paleogeographic map showing major depositional systems that existed on the shelf 

120 ka to 22 ka (stages 5e–3 highstand). 

The different lobes of individual deltas are numbered in chronological order. RGD = Rio Grande Delta, CD 

= Colorado Delta, BD = Brazos Delta, WLD = Western Louisiana Delta, LD = Lagniappe Delta, EMD = 

Eastern Mobile Delta, WFLAD = West Florida–Alabama Delta, and AD = Apalachicola Delta. Reprinted 

from Anderson et al. (2004)

Anderson et al. (2016) divided the channel belts of Berryhill (1987) and Suter (1987) into two separate 

drainage systems. They recognized an eastern channel belt complex, which they inferred to be a paleoMississippi River channel belt complex that is characterized by somewhat wider, more closely spaced 

channel belts. This eastern set of channel belts displays lateral accretion, generally less than a kilometer, 

indicating modest channel sinuosity, and occupied an area at least 150-km wide. It is attached to an 

eastern shelf margin delta, which Suter and Berryhill (1985) called the “Mississippi Delta”. This delta has 

not been directly dated but is assumed by Anderson et al. (2016) to be an MIS 3 feature because the 

Mississippi River is known to have fed into the Mississippi Canyon during MIS 2. 

As defined by Anderson et al. (2016), the western channel belt complex consists of channel belts that 

converge seaward and occupy an area about 80 km wide. Individual channel belts are more than 35 m 

deep, and have width-to-depth ratios generally greater than 30:1 (Suter 1987). It is inferred that these 

channel belts lay on the surface of the coastal plain without any associated paleovalleys because these 

channel belts are single, not multi-story. One of the channel belts appears to be an offshore extension of 

one branch of the paleo-Mississippi Lafayette Meander Belt and associated Avoyelles alloformation, 

which ends in a large shelf phase and shelf margin delta, known as the “Western Louisiana Delta.” This 

delta has been radiocarbon-dated by Wellner (2004) at 30 ka BP. However, radiocarbon dates in that age 

range have sometimes proved unreliable and misleading within the Gulf of Mexico, as noted by Stapor 

and Tanner (1973), and should be corroborated using optically stimulated luminescence dating.
27

3.3.4 Last Glacial Maximum sea level MIS 2–LGM

During the last glacial cycle, the maximum lowstand of sea level occurred during the Last Glacial 

Maximum (MIS 2), between 30 and 19 ka BP. Sedimentological and age data from the Huron Peninsula, 

New Guinea, indicated that sea level fall to this lowstand began about 32 ka with a rapid drop of 30 to 50 

m over a period of 1 to 2 ka (Lambeck et al. 2002). The exact amount of global sea level fall is 

controversial, but consensus based on ice-volume models and the Barbados coral record is near -125±5 m

(Nakada and Lambeck 1989, Fairbanks 1989, Bard et al. 1990, Fleming et al. 1998).

The elevation of the relative sea level lowstand in the Gulf of Mexico during MIS 2 is still poorly 

constrained. Using glacial isostatic adjustment modeling, Törnqvist et al. (2006) estimated that relative 

sea level at this time was 34 m to 40 m “below shelf edge.” Bart and Ghoshal (2003) concluded, from an 

examination of the shelf-margin delta geomorpology within the northeastern Gulf of Mexico, that relative 

sea level during the MIS 2 was about -65 m to -85 m depending on location along the shelf edge. Along 

the southeastern Louisiana continental shelf, Roberts et al. (2004) argued that relative sea level was at 

between -85 m and -89 m during the formation of the Lagniappe delta about 19 ka. They also reported a 

still inexplicable drop in relative sea level of sea level of 30 m over a 1,000 to 3,000 period about 16–18 

ka. The ICE5G_120550_Full model of Martini (2015) produces relative sea level values similar to those 

reported by Bart and Ghoshal (2003) and for 19 ka by Roberts et al. (2004).

On the basis of limited data, it appears that the paleoshoreline in the northern Gulf was situated at or near 

the shelf break (approximately -65 m to -85 m water depth) during the MIS 2 glacial lowstand of sea level 

(Bart and Ghoshal 2003, Roberts et al. 2004). As a result, almost the entire continental shelf was a 

subaerially exposed coastal plain and available for occupation by humans. Prominent features of the 

lowstand coastal plain included fluvial valleys and lowstand, shelf margin deltas. Offshore, they fed large 

amounts of sediment to delta-fan complexes, slope fans, and sediment-gravity-flow deposits, along with 

hemipelagic drapes and contourites on the continental slope (Anderson et al. 2004, 2014, and 2016). The 

exact duration of time for which this lowstand coastal plain was subaerially exposed is still uncertain, but 

it lasted only a few thousand years. This surface, buried by younger transgressive deposits, forms a 

prominent bounding unconformity known by a variety of names including “Base of Holocene, Late 

Wisconsinan unconformity,” and “Pleistocene-Holocene surface.” Simms et al. (2007) have constructed a 

paleogeographic map of this surface (Figure 3.5).
28

Figure 3.5. Paleogeographic map showing the major depositional systems that existed on the 

shelf and upper slope during the 22 ka to 16 ka (MIS 2) lowstand. 

Lowstand paleovalleys are labeled as follows. RGV = Rio Grande, CV = Colorado, BV = Brazos, T/SV = 

Trinity–Sabine, WMV = west Mobile, and EMV = east Mobile. MC = Mississippi Canyon. Reprinted from 

Anderson et al. (2004)

During the MIS 2 sea level lowstand, fluvial systems within the northeastern Gulf of Mexico coastal plain 

extended their valleys to the shelf edge and incised their valleys deeply into the surface of the coastal 

plain. Because of the low gradients of the east Texas and western Louisiana continental shelves, the 

lowstand paleovalleys of the Calcasieu, Mermentau, Sabine, and Trinity rivers become broader and 

shallower in a gulfward direction (Figure 3.6; Anderson et al. 2016). Typically, these paleovalleys are 40 

m deep at the present shoreline (Blum and Price 1998). Within the paleovalleys of the Sabine and Trinity 

rivers, well-defined fluvial terraces have been found (Thomas and Anderson 1994, Blum et al. 1995,

Rodriguez et al. 2005). The presence of similar terraces within the onshore valleys of the Calcasieu and 

Mermentau rivers indicates that, depending on transgressive shoreface and tidal ravinement depths,

terraces likely are preserved in the offshore paleovalleys of these rivers. They also suggest that the 

incision of the onshore valleys, and presumably the equivalent offshore paleovalleys of the Calcasieu, 

Mermentau, Sabine, and Trinity initially occurred during either MIS 3 or even MIS 4.

3.3.5 Holocene Sea Level: MIS 1

Starting about 19 to 18 ka and lasting until about 6 ka, rapid deglaciation of continental ice sheets 

occurred. As a result, at the end of the Last Glacial Maximum, about 19 ka, global sea level 

correspondingly rose rapidly by around 125±5 m to near present about 6 ka. After about 6 ka, the 

surviving ice sheets have stayed relatively stable and, as a result, rates of sea-level rise decelerated 

significantly to an average rate of 1.4 mm/yr (Lambeck et al. 2002).

However, it is well documented that global sea level during the last deglaciation did not rise continuously, 

but in a series of steps with different rates. There was a distinct and short period of stability during the 

Younger Dryas, between 12.5 and 11.5 ka, when ice melt briefly ceased (Chappell and Polach 1991,

Lambeck et al. 2002). Also, episodes of extremely rapid global sea level rise that are related to 
29

accelerated ice-sheet decay and melt-water discharge have also been recognized. The two most widely 

recognized and accepted are referred to as melt-water pulse 1A (MWP-1A) and melt-water pulse 1B 

(MWP-1B) (Fairbanks 1989; Bard et al. 1990). Although the precise timing, duration, amplitude, and 

rates of global sea-level rise associated with these events remain a subject of continuing discussion, 

MWP-1A is generally accepted as having started at 14.5 ka, after which global sea level rose by 40–

55mm/ yr for about 500 years (Edwards et al. 1993). In the case of the more ambiguous and less certain 

MWP-1B, it is speculated that it began about 11.5 ka, lasted for about 200 years, and produced global 

rates of sea-level rise comparable to those of MWP-1A (Fairbanks 1989). Additional melt-water pulses 

have been reported as beginning at 19 ka (Clark et al. 2004) and 7.6 ka (Blanchon and Shaw 1995);

however, their significance and extent have been disputed (Woodroffe and Horton 2005).

As discussed by Anderson et al. (2008) and Anderson et al. (2016), the relative sea level history of the 

northeastern Gulf of Mexico is well constrained for the terminal Pleistocene (latter part of MIS 2), and the 

Holocene (MIS 1). Starting about 19 ka the rate of sea level rise is unconstrained, but was rapid and 

punctuated with intervals of very rapid sea level rise until about 10 ka. Between about 10ka and 7.5 ka, 

the mean rate of sea level rise was about 4.2 mm/yr. At about 7.5 ka, the mean rate of sea-level rise 

slowed to 1.4 mm/yr until about 4.0 ka. At that time, the current highstand began when the rate of sealevel rise slowed to ~0.4 to 0.6 mm/yr. It was at about 4±1 ka when most of the current strandplains, 

barrier islands, peninsulas, and chenier plains began to form (Anderson et al. 2014 and 2016, Milliken et 

al. 2008a). 

During the late MIS2 and MIS 1 marine transgression, preservation of the coastal deposits was minimal

(e.g., Anderson et al. 2016; Rodriguez et al. 2004; Penland et al. 1988; Belknap and Kraft 1981). Some

fluvial deposits, such as channels, floodplains, and terraces and associated bayhead delta, estuarine, and 

tidal deposits are preserved and comprise the fill within major paleovalleys, such as the Sabine-Trinity 

system (e.g., Thomas and Anderson 1994, Rodriguez et al. 2005, Simms et al. 2006). As will be discussed 

in more detail, low preservation potential outside of the major paleovalleys is the result of deep shoreface 

and tidal ravinement processes (Goff et al. 2015, McBride et al. 2004, Penland et al. 1988). As a result, 

the record of paleosurfaces and coastal evolution on the continental shelf remains fragmentary outside of 

paleovalleys (Nelson and Bray 1970, Thomas and Anderson 1994, Anderson et al. 2014 and 2016, Simms 

et al. 2006). 
30

4. Shelf Morpho-sedimentary Response to Sea Level Change

A variety of processes act to modify paleolandscapes of a coastal plain as changes in relative sea level and 

sediment supply result in either the landward movement of the shoreline, a transgression, or the seaward

movement of the shoreline, a regression. As indicated by the previously discussed sea level history of the 

last glacial–interglacial cycle, MIS 5 through 1, changes in shoreline can be subdivided into four general 

modes of coastal plain and continental shelf response to relative sea level change: 1) continuous 

transgression, 2) episodic transgression, 3) normal regression, and 4) forced regression. 

4.1 Transgressive Surfaces

The significant modification of a subaerial coastal plain occurs during both continuous and episodic 

transgression as the result of relative sea level rise and resulting formation of transgressive surfaces. The 

surface formed is either a flooding or ravinement surface. A flooding surface is a disconformity produced 

by transgression that marks an abrupt increase in water depth. For example, this surface can be the abrupt 

upward change from either the subaerial exposed surface of a coastal plain to marsh and eventually 

lagoonal sediments or from shallow-water sediments into those that accumulated in deeper water. If this 

surface exhibits significant erosion, it is known as a ravinement surface. A ravinement surface is an

erosional surface produced in the coastal zone, by either wave erosion or by tidal scour in channels within 

estuaries, during marine transgression, often of a formerly subaerial environment. An in-depth review of

transgressive surfaces can be found in Cattaneo and Steel (2003).

4.1.1 Coastline flooding (ravinement) surfaces

Of significance to the preservation of paleolandscapes and archaeological sites is what is designated in 

this report as a coastal flooding surface. A coastal flooding surface is a disconformity produced by 

transgression that marks an abrupt change from the terrestrial sediments of a coastal plain to either 

lagoonal or marine sediments. This disconformity can be either the relatively intact surface of a coastwise terrace, fluvial terrace, floodplain, or delta plain or, more commonly, eroded to form a minor 

ravinement surface.

Based on their environment of formation, three types of coastal flooding surfaces are end members of a 

continuum that grade into each other and form a single, regional coastal flooding surface. They are: 1)

bayline flooding (ravinement) surface; 2) lagoonal flooding surface; and 3) soundline flooding 

(ravinement) surface. The bayline flooding surface is the disconformity produced by transgression of a 

bay over floodplains, bayhead deltas, and terraces within either an estuary or ria. Depending on the size of 

the estuary involved, there may not be enough fetch for significant wave action and the resulting erosion 

and ravinement development to take place. In addition, the accumulation of marsh or swamp deposits also 

can bury the valley floor and associated terraces under a protective layer of sediment. A lagoonal flooding 

surface is the disconformity produced by transgression of a lagoon over a coastal plain consisting of 

coast-wise terraces and non-incised fluvial channel-belts. In case of a narrow lagoon behind barrier 

islands that would characterize such a coast compared with a wider sound, the fetch would be narrow 

enough to limit wave erosion and marsh accumulation along the landward edge of a lagoon and would 

offer some protection to the surface of the subsiding coastal plain and associated channel belts. Finally, a 

soundline flooding surface is the disconformity produced by transgression of a sound over either a coastal 

plain or delta plain. The fetch of a sound would be large enough to allow significant, but still limited,

wave erosion. As a result, a soundline flooding surface typically is a minor ravinement surface with 

erosion on a magnitude of a meter or so. An example of a soundline flooding surface is the eroded surface 

of the St. Bernard delta lobe underlying Chandeleur Sound between the Chandeleur Islands and west to 

the eroding edge of the subsiding Mississippi Delta Plain.
31

4.1.2 Shoreface ravinement surface

A major transgressive surface is the shoreface ravinement surface. The shoreface ravinement surface is a 

well-defined erosion surface produced in the coastal zone, by wave erosion within the shoreface during 

marine transgression along the seaward edge of either a barrier island chain or coastal plain shoreline. As 

the shoreline moves landward, the shoreface truncates back-barrier coastal sediments and underlying 

coastal plain sediments to depths of 8 to 10 m before the wave base is exceeded (Wallace et al. 2010). If 

the rate of sea level rise is relatively low, the depth of erosion is deep enough that it typically erodes 

through the coastal flooding surface and the marsh, lagoonal, beach, and other back-barrier coastal 

sediments that lie landward of the shoreface and into the underlying coastal plain (Figure 4.1). If the rate 

of relative sea level rise and sedimentation is relatively rapid, the shoreface might not be deep enough to 

reach the coastal flooding surface and allow preservation of it and some back-barrier coastal sediments 

overlying the coastal flooding surface (Figure 4.2) (Belknap and Kraft 1981 and 1985, Swift and Thorne 

1991, Thorne and Swift 1991).

Figure 4.1. Marine transgression by shoreface retreat in which shoreface erosion associated with 

a rise in sea level removes older coastal deposits, except within the paleovalley, and 

lagoonal flooding surface and creates marine ravinement surface. 

The marine ravinement surface truncates older coastal plain surface and destroys any associated cultural 

deposits lying on it. The high-tide level [HT] and the low-tide level [LT] are shown at times 1 and 2. 

Modified after Waters (1992). 
32

Figure 4.2. Marine transgression by in-place drowning during period of rapid relative sea level 

rise. 

The shoreline jumps landward as relative sea level rises rapidly from a lower position to a higher position. 

Although a large block of coastal sediments and a segment of lagoonal flooding surface are preserved 

because the surf and swash zones have little time to erode the older barrier island sequence, the surface 

of the older barrier island is reworked by surf and swash zones and tidal currents. Modified after Waters 

(1992).

The shoreface ravinement surface is created by waves hitting the shoreface and interacting with 

longshore, geostrophic, and tidal currents. Waves are regarded as an overwhelmingly destructive force in 

terms of terrestrial paleosurfaces and associated archeological resources. The depth of erosion by 

shoreface erosion varies greatly. Most commonly, it appears to be within the 8 m to 10 m range (Belknap 

and Kraft 1981 and 1985).

The formation of a shoreface ravinement surface by wave action and associated longshore currents has 

important ramifications for the preservation of paleosurfaces and associated archaeological deposits. First, 

the surfaces and sediments of barrier, e.g., beach ridges, dunes, and beaches, and back-barrier 

environments, are typically prone to being destroyed by shoreface erosion, except under the most unusual 

circumstances. Beach ridges and dunes, unless they consisted of either carbonates or mixed carbonateclastic sediments cemented into beachrock, are typically destroyed by shoreface erosion during 

submergence. Second, the barrier and backbarrier sediments, coastal flooding-ravinement surface, and the 

underlying Pleistocene sediments are vulnerable to erosion depending on the depth of erosion. In general, 

the slower relative sea level rises from a combination of subsidence and eustatic sea level rise, the longer 
33

the shoreface is subject to the erosive power of waves and the deeper it cuts into the barrier and 

backbarrier sediments. If the rate of relative sea level rise is slow enough, the shoreface eventually erodes 

through the barrier and backbarrier sediments and coastal flooding-ravinement surface and into the 

underlying Pleistocene sediments. The faster relative sea level rises, the shorter the shoreface is subject to 

the erosive power of waves and it might not erode completely through the backbarrier sediments and 

leave the coastal flooding-ravinement surface and any underlying deposits intact. Finally, the lighter 

materials found within any barrier, backbarrier, and underlying sediments that are eroded will be 

transported either in the direction of littoral drift or shoreward onto the beach. Denser and larger material, 

e.g., shells, artifacts, and pedogenic nodules, will tend to remain in place, as the size and weight for 

materials that can be transportable by wave action and longshore currents are limited. This process will 

concentrate these clasts, including artifacts, as an erosional lag associated with and overlying the 

shoreface ravinement surface (Belknap and Kraft 1981 and 1985, Murphy 1990).

Geostrophic currents are a dominant process on the Texas-Louisiana continental shelf. They are currents 

that are sustained by a balance between a pressure gradient force, e.g., either the Gulf Stream or hurricane 

storm surge, and the Coriolis force. Geostrophic currents are insufficient as an erosional force in the surf 

zone. However, they occur in combination with combined flow to transport sand offshore on the shelf 

during storms. They will have the same general effects as wave action on shoreface erosion and 

associated clast (artifact) transport and concentration (Snedden 1985, Snedden et al. 1988, Siringan and 

Anderson 1994).

4.1.3 Tidal ravinement surfaces

Another transgressive surface that is often important in terms of the preservation and destruction of 

paleosurfaces is the tidal ravinement surface. In a wave- and tide-dominated environment, the geometry 

and extent of a tidal ravinement surface is a function of the tidal ravinement processes, which characterize

either an estuary or tidal inlet along a barrier shoreline. 

The tidal inlets associated with barrier island chains of a subsiding delta lobe can be more or less 

anchored by the paleochannels of deltaic distributaries (Kulp et al. 2007). In case of such tidal inlets, 

which Féniès et al. (2010) refers to as a “anchored tidal ravinement,” the inlet remains in place and the 

cross-section of the original distributary channel is enlarged with time as the bay behind it and associated 

tidal prism enlarges. This has the effect of eroding and eventually destroying the original channel and the 

natural levees on either side of a deltaic distributary’s paleochannel (Boyd et al. 2006). 

Other tidal inlets are associated with barrier islands within broad estuaries; the tidal inlet often migrates in 

response to lateral spit accretion. These processes create a laterally persistent, coast-wise tidal ravinement 

surface (Miner et al. 2007). Such tidal ravinement surfaces often cut into, across, and cap the fills of 

paleovalleys and adjacent interfluves along former shoreline positions. Depending on the depth and lateral 

extent of the paleovalley, all or part of the fills can be removed by tidal ravinement. As a result, the 

preservation of the bayline (coastline) flooding surface, terrace surfaces, and other paleosurfaces will 

depend on the depth of erosion by the tidal surface. All else being equal, the thickness of estuarine and 

fluvial deposits, which is a function of the size and depth of the paleovalley and size of the drainage 

basin, will be a major factor (Boyd et al. 2006, Féniès et al. 2010).
34

4.2 Paleovalleys and Channel Belts

As previously noted, formation of paleovalleys as the result of forced regression created by glacial and 

stadial sea level fall and lowstand has repeatedly occurred within the northwestern Gulf of Mexico. 

During the subsequent interstadial or interglacial sea level rise and highstand, these paleovalleys started to 

fill with sediments before the subsequent migration of the shoreline over them. This creates an 

aggradational package of sediments that often has a significant potential for the preservation of 

paleosurfaces (Anderson et al. 2004, 2014, and 2016, Suter 1987, Suter et al. 1987).

4.2.1 Forced regression, paleovalleys, and channel belts

Falling relative sea level causes basinward retreat of the shoreline, a process termed “forced regression”

by Plint (1996). The fall in relative sea level that results from a forced regression exposes the continental

shelf as coastal plain to subaerial erosion. Within the northwestern Gulf of Mexico where the gradient of 

the continental shelf is nearly equal to or greater than the gradient of its river course, as is typical of 

major, extrabasinal river systems, a river valley will vertically incise (downcutting) as its mouth migrates 

seaward. Depending on the degree of relative sea level fall and gradient of the continental shelf, its course 

continually grades downward to progressively lower sea levels and creates a deeply incised valley on the 

exposed shelf (Anderson et al. 2004, 2014, and 2016, Suter 1987, Suter et al. 1987). 

If the forced regression episodically alternates with periods of either stability or sea level rise, then 

channel belts will likely be created that fill the valley at each period by lateral migration of the 

contemporaneous channel. At the same time lateral planation by river scour creates an erosional surface 

that is buried as soon as it is created. This is called a “regional composite scour surface” and forms the 

bottom of a valley. The resumption of relative sea level fall will typically result in incision of the former 

channel belt and associated floodplain and formation of subaerial paleosurfaces in the form of terraces 

lying above and flanking a newly developed floodplain and channel belt. If the fluvial incision is deep 

enough, the lateral planantion of the new channel belt will cut through and destroy both the older fluvial 

deposits and its composite scour surface. As a result, the bottom of the valley will be a composite of 

regional composite scour surfaces that never existed as a single subaerial topographic surface and that 

formed diachronously at the channel-belt scale over the entire fall to rise of a base-level cycle (Holbrook 

and Bhattacharya 2012).

As is typically the case of intrabasinal river systems with significantly smaller drainage basins and river 

gradients that are largely equal to those of the continental shelf exposed during the forced regression, 

either no or relatively insignificant entrenchment will occur (Anderson et al. 2004, 2014, and 2016, Suter 

1987, Suter et al. 1987). Instead, the river will form a channel belt that is generally level with or occupies 

a fluvial ridge lying above the level of the coastal plain as observed for the surface of the MIS 5 coastparallel terraces in East Texas by Aronow (1971) and Van Siclen (1985 and 1991).

In case of Southwest Louisiana, an incised, post-MIS 5e Red River paleochannel has been observed near 

Lake Charles, Louisiana in interpreted LiDAR DEMs. In this case, it appears that the lowering of relative 

sea level during the latter part of MIS 5 was at a rate that allowed for the downcutting of a paleochannels, 

but prohibited the development of a paleovalley by lateral migration. The entrenchment of this channel 

appears to have beheaded what are likely fluvial or deltaic distributary ridges as it extended its mouth 

gulfward.
35

Modern valleys are defined as elongate topographic lows that contain rivers, which lack the discharge to 

routinely flood beyond the valley incision and onto the interfluves. However, this criterion often cannot 

be applied effectively in identifying buried, prehistoric paleovalleys or differentiating them from 

unincised channel belts that once were part of the relatively flat surface of submerged coastal plain. To 

differentiate a channel belt lying at the bottom of a paleovalley and a channel belt lying unincised on the 

surface of a coastal plain, a proxy for incision associated with a channel belt is required (Holbrook and 

Bhattacharya 2012). Thus, the recognition of a paleovalley requires either (a) the preservation of 

recognizable valley walls that extend above the level of the floodplain or (b) a sand body lying beneath a 

ravinement surface that consists of two or more sets (stories) of fluvial sands. In the case of single-story

channel belts with upper surfaces truncated by regional ravinement surfaces, differentiating between 

coastal plain channel belts and deeply eroded paleovalleys might be difficult, if not impossible.

4.2.2 Transgressions, paleovalleys, and channel belts 

The preservation potential of a channel belt during a marine transgression varies greatly depending on its 

position relative to the surface of a coastal plain. The paleosurfaces of channel belts that either lie on the 

surface of a coastal plain or are part of a fluvial ridge are the least likely to survive a marine transgression,

because they will be directly exposed to shoreline erosion as the coastal plain is submerged. In contrast, 

both lying below the level of shoreface erosion and the transgressive backfilling of paleovalleys with a 

prism of sediments often protects the surface of channel belts, their associated terraces, and other 

paleosurfaces during a marine transgression (Pearson et al. 1986). 

As described by Blum and Törnqvist (2000), a rise in relative sea level results in channel shortening, 

decreases in the distance over which sediments can be stored and, in most cases, a flattening of the 

channel slope. Even though discharge is conserved, these changes result in corresponding downstream 

decreases in stream power, sediment transport rates, and, ultimately in net aggradation. As a result, rising 

sea level causes a prism of initially fluvial and later estuarine sediments to accumulate within a valley 

system that extends and tapers upstream from the contemporaneous shoreline. As relative sea level rises, 

this sedimentary prism buries the surfaces of channel belts, floodplain, and in some cases even low-lying 

terraces (Blum and Törnqvist 2000, Shen et al. 2012, Blum et al. 2013).

During the periodic episodes of abrupt relative sea level rise, rapid flooding of an estuary will occur. Such 

episodes cause the overstepping and flooding of bayhead deltas and segments of fluvial terraces and 

landward shift in estuarine facies and the formation of flooding surfaces. If sea level rise occurs fast 

enough, the surface of bayhead deltas and terraces are buried and preserved as paleosurfaces within the 

fill of a paleovalley (Rodriguez et al. 2005, Milliken et al. 2008b and 2008c). If sedimentation rates are 

high and the paleovalley deep enough, these paleosurfaces can survive subsequent destruction by 

shoreface erosion and formation of a ravinement surface (Pearson et al. 1986).

As seen in aerial photographs and LiDAR DEMs, the surface of the MIS 5 coastal-parallel terrace of 

southwest Louisiana and southeast Texas is characterized by numerous relict Pleistocene channel belts 

and fluvial and distributary ridges along with younger Holocene drainages (Fisk 1948, Bernard and 

LeBlanc 1965, Aronow 1971 and 2004, Van Siclen 1985 and 1991, Heinrich 2007). During a marine 

transgression the erosion of the upper 8 m to 10 m of the MIS 5 coastal-parallel terrace would destroy 

distributary and fluvial ridges and constructional surfaces of relict channel belts. At best, only the 

truncated coarse-grained basal channel fills of the larger intrabasinal rivers and streams would remain. 

Although, the surface of the MIS 5 coastal-parallel terrace might survive the initial flooding by 

backbarrier lagoons and burial by lagoonal sediments, it would be eventually destroyed from shoreface 

erosion and formation of a ravinement surface.
36

4.3 Deltas and Chenier Plain

During either stillstands or generally low rates of relative sea level rise, segments of the northwestern 

Gulf of Mexico coastal plain have experienced extensive periods of normal regression (Figure 4.3). In this 

part of the Gulf of Mexico, normal regression is the seaward advance of the coastline as a result of the 

progressive addition of sediment to either the front of a shoreline, as in case of the Louisiana Chenier 

Plain, or a delta system. In the case of major delta systems, normal regression is typically associated with 

aggradation of the delta plain because of subsidence-induced relative sea level rise. When delta lobes are 

abandoned because of delta switching, deltas are locally and periodically subject to marine transgression 

(Coleman and Roberts 1998a and 1998b).

Figure 4.3 Normal regression in which coastline progrades gulfward in response to sediment 

accumulation during period of constant or slowly rising relative sea level. 

The creation of new and productive delta or chenier plains may promote the preservation of cultural 

resources.

4.3.1 Regressive and transgressive phases of deltas

Although deltas like the Mississippi River Delta have, until historic times, been actively prograding 

outward into the Gulf of Mexico, each delta lobe evolves through a cycle of events that begins with 

sedimentation and rapid growth as part of a constructional or regressive phase. The constructional or 

regressive phase of delta growth typically consists of the rapid infilling of lakes and bays, followed by the 

local gulfward progradation of the coastline and the deposition of a thick wedge of terrigenous clastic 

sediment on the adjacent shelf. During this phase, progradation creates a relatively thin shelf delta with 

numerous distributary channels, merged distributary mouth bar sands, and distributary channels that cut 

below the base of delta deposit (Penland et al. 1988, Roberts 1997, Coleman and Roberts 1988a and 1988

b).
37

The constructional or regressive phase is followed by the abandonment of a delta lobe by its distributary 

channel, which cuts off water and sediment flowing to it, and results in the deltaic deterioration of its 

destructional or transgressive phase. During this phase, distributaries of the previously active delta lobe 

become erosional headlands as marine transgression and associated reworking results in the landward 

migration of the shoreline. Sandy sediments eroded from distributaries that form the headlands is 

reworked laterally by waves to form flanking barrier islands that enclose restricted interdistributary bays. 

As the deltaic plain subsides below sea level, the transgressive barrier island arc is separated from the 

retreating mainland shoreline and eventually forms an inner-shelf shoal. After submergence, the innershelf shoal continues to be reworked into a marine sand body on the inner continental shelf. In addition, 

organic sediments would continue to accumulate in the interior and distal portions of the delta plain 

relative to the shoreline. This process would bury natural levees and other surfaces within these portions 

of the delta plain until submerged by relative sea level rise (Penland et al. 1988, Roberts 1997, Coleman

and Roberts 1988a and 1988b).

4.3.2 Regressive and transgressive phases of Louisiana Chenier Plain

The Louisiana Chenier Plain within Southwest Louisiana consists of multiple, shore-parallel, sand-rich 

ridges that range in elevation from 2 m to 6 m. They rest on and are separated from one another by finergrained, clay-rich sediment. They evolved during the Holocene as the result of the periodic progradation 

of mudflats that were interrupted by regressional episodes that eroded and reworked existing mudflats and 

concentrated the coarse sediment fraction into sandy or shelly ridges. The mudflats reflect the 

accumulation of fine-grained sediment discharged by the Mississippi River and transported westward. 

During episodes of non-deposition due to eastward shifts of the Mississippi River deposition, coarsegrained sediment and shells are concentrated along the shore forming chenier ridges. Renewed mudflat 

deposition and isolation of the chenier ridges occurs when a new supply of sediment from the Mississippi 

River distributaries is again transported westward. Currently, mudflats along segments of the Louisiana 

Chenier Plain are presently prograding from sediments being transported westward from the Atchafalaya 

River (Gould and McFarlan 1959, Penland and Suter 1989).

Because of their low relief, the surfaces of both the Mississippi River Delta and the Louisiana Chenier 

Plain, like the beach ridges of barrier islands and delta plains, would be susceptible to rapid submergence 

and in-place drowning from a period of rapid, episodic sea level rise. In-place drowning would cause an 

abrupt lateral shift of the shoreline, called “overstepping,” from its current shoreline to the boundary 

between these plains and the Pleistocene coast-parallel terraces. Although this would likely prevent 

complete destruction from shoreface erosion of the bulk of the deposits underlying these plains, their 

surfaces would still be significantly disturbed and, in case of the cheniers, reworked into inner-shelf 

shoals similar to the proposed origins for Herald and Sabine banks.

4.3.3 Forced regression, deltas, and beach ridges

During the last glacial-interglacial cycle, forced regression impacted delta and beach ridge systems before

the human occupation of the study region during MIS 5b through MIS 3 (Figure 4.4). Forced regressions 

have resulted in abandonment and stranding of the relict Houston Ridge segment of the Ingleside 

Paleoshoreline within Calcasieu Parish, Louisiana, and the beach ridges, including Milton Island, of the 

Ponchatoula beach system adjacent to the north shore of Lake Pontchartrain in southeast Louisiana. Also, 

off-lapping shelf deltas of the West Louisiana Delta within the southwest Louisiana continental shelf, are 

attributed to forced regression (Figure 3.4).
38

Figure 4.4. Forced regression in which coastline moves gulfward in response to relative sea level 

fall. 

Falling sea level allows for the preservation of cultural resources and opening of new land for occupation.
39

5 Preservation of Paleolandscapes and Cultural Resources on the 

Northern Gulf of Mexico Continental Shelf 

5.1 Landscape: Cultural Resources Associations

A number of studies–including McIntire (1958, 1971), Coastal Environments, Inc. (1977), Gagliano et al. 

(1982), Gagliano (1984), Goodwin et al. (1991), Weinstein et al. (1992), Dunbar et al. (1991), Faught and 

Donoghue (1997), and Faught (2004)–on the northern Gulf of Mexico plain have documented a very 

close relationship between the pattern of prehistoric settlement and modern geomorphic surfaces and their 

associated landforms. Within the northwestern Gulf of Mexico coastal plain, these geomorphic surfaces 

can subdivided into the Mississippi Delta Plain; incised fluvial valleys of the Sabine, Trinity, and other 

rivers; the chenier plain, barrier islands, salt dome “islands;” and upland coast-parallel terraces. In the 

northeastern Gulf of Mexico Coastal Plain, another geomorphic terrain, the karst uplands dominates large 

parts of the landscape.

The Mississippi Delta Plain was widely used by prehistoric cultures as documented by a number of 

studies. The numerous archaeological studies and cultural resource management surveys have shown that 

archaeological sites have specific associations with various deltaic landforms. Deltaic archaeological sites 

are largely associated with the crests of major and minor natural levees. Specifically, preferred locations 

include crests of natural levees along trunk channels; natural levees of crevasse distributaries near swampmarsh interfaces; on distributary levees at juncture with crevasse subdeltas; on distributary levees at the 

heads of major subdelta lobes; and natural levees near the mouths of active distributaries (Figure. 5.1). 

Other preferred locations for cultural deposits associated with deltas include the shores of bays and lakes 

within deltas and the beach ridges of barrier arcs and distributary mouth accretion ridges. 

Figure 5.1. Preferred prehistoric site locations in the Mississippi Deltaic Plain. 

Reprinted after Gagliano et al. (1982).

Within deltas, the largest accumulations of cultural deposits are typically associated with the crests of 

natural levees typically either along the cutbank side of meander loop cutoffs, at junctions of tributaries 
40

and distributaries, or along crevasse distributaries. Most commonly, the cultural deposits lie on the 

surface of the natural levee and, thus, accumulated after the distributary had been abandoned as an active 

course of the parent river system. This is typical of the major cultural deposits associated with intensive 

occupation areas, burials, and so forth, found within the Mississippi River Delta (Coastal Environments 

Inc. 1977, Gagliano et al. 1982, Weinstein and Gagliano 1985, Weinstein et al. 1992). Gagliano (1984) 

noted that cultural deposits also have been found interbedded with sediments of a natural levee, indicating 

that the site was occupied while the distributary paleochannel was active and within the adjacent channel 

fill that accumulated after the distributary had been abandoned (Figure 5.2). Gagliano (1984) briefly noted 

that cultural materials might be found interbedded with backswamp clays and peats where they 

accumulated either while a distributary was active or after it had been abandoned. Such cultural deposits 

are usually deposited in standing water and may contain well-preserved perishable artifacts and ecofacts.

Figure 5.2. Cross section of terrace of “High” Deweyville at McNeill Ranch Site, Victoria County, 

Texas. 

Modified from Aiuvalasit (2007).

As discussed by Coastal Environments, Inc. (1977) and Gagliano et al. (1982), the beach ridges of the 

Louisiana Chenier Plain, Texas barrier islands, barrier arcs of subsiding delta lobes, and distributary 

mouth accretion ridges are visually prominent, well-drained, and suited landforms for prehistoric 

habitation. As a result, significant accumulations of cultural deposits are commonly associated with them. 

A prime location for the accumulation of cultural deposits is the ends of beach ridges either at the mouth 

of chenier plain bayous or within barrier-spits adjacent to tidal channels. Because of the nature of the 

depositional environments, cultural deposits are typically surficial in nature and of restricted thickness 

(Coastal Environments, Inc. 1977, Gagliano et al. 1982, Gagliano 1984).
41

The distribution of archaeological sites within fluvial valleys of the major rivers within Southeast Texas 

and the Louisiana Coastal Plain has been studied by Gagliano et al. (1982), Pearson et al. (1986), Abbott 

(2001), and Aiuvalasit (2007). They found that, as within deltas, major sites can be found on the crests of 

natural levees found within active floodplains. Based on research within the Mississippi River alluvial 

valley, Weinstein (1981) argues that prehistoric cultures preferentially occupied the natural levees 

bordering abandoned channel segments, with the largest sites being located adjacent to oxbow lakes 

within these segments, and avoided the natural levees bordering active channels. However, it is important 

to note that Gagliano et al. (1982), Pearson et al. (1986), Abbott (2001), and Aiuvalasit (2007) conclude 

that for incised fluvial valley systems, such as the Calcasieu-Sabine-Trinity, the margins of terraces and 

valley walls nearest active channels were the preferred location for prehistoric occupations and 

accumulation of significant cultural deposits, and not the natural levees, as observed in the delta plain. In 

southeast Texas, Abbott (2001) and Aiuvalasit (2007) found that the margins of terraces were also the 

location for the accumulation of wind-blown sediment periodically during the Holocene. The periodic 

accumulations of eolian sediments lead to the creation of well-stratified cultural deposits (Figure 5.2). 

During the flooding of alluvial valleys by sea level rise, the margins of valley walls and terraces 

overlooking an estuary also were the locations of prehistoric occupations and the accumulation of 

significant cultural deposits (Figure 5.3) (Gagliano et al. 1982).

Figure 5.3. Location of shell middens and shell content in a hypothetical coastal estuary. 

Also, changes in shell content of a hypothetical shell midden resulting from sea level fluctuations. 

Reversal in apparent salinity within the stratified site results from a rise in sea level (A–C), followed by 

stream progradation (D–E). A: oldest occupation; E: youngest occupation. Reprinted from Coastal 

Environments, Inc. (1977).
42

Although much less well-studied than other coastal plain surfaces, interfluves (uplands separating valleys) 

of Pleistocene coast-parallel terraces have distinct landform-associated cultural deposits. When sea level 

was much lower than present during the last glacial-interglacial cycle, these uplands extended out onto the 

continental shelf and exhibited a collection of relict natural levees of channel belts and deltaic 

distributaries and relict beach ridges. Better drained because of their coarser texture and greater elevation, 

the crests of these features are likely the preferred location of prehistoric occupation and accumulation of 

cultural deposits, as suggested by limited surveys conducted for cultural resource management studies,

such as Weinstein et al. (1979a). Within the coast-parallel terraces, cultural deposits also have been found 

associated with the summit of innumerable pimple mounds that cover the surface of the Pleistocene coastparallel terraces (Aten and Bollich 1981, Jones and Schuman 1988). As discussed further below, these 

uplands on the exposed shelf have little to no preservation potential during sea-level rise. 

5.2. Paleolandscape and Cultural Resources Preservation 

From the previous discussion, it is obvious that large areas of the Gulf of Mexico continental shelf were 

exposed as land during the last glacial-interglacial cycle, but have become submerged and reworked as a 

result of sea level change. It is also well documented that human populations were present within the Gulf 

of Mexico Coastal Plain during at least the last 13,000 years and possibly for the last 20,000 years. This 

means that the Louisiana and East Texas continental shelves were potentially the sites of pre-historic 

human occupation. The main limiting factor in terms of preservation potential is whether the 

paleosurfaces on which cultural deposits might have accumulated survived the Late Wisconsin–Holocene 

transgression.

5.2.1 Transgression: coastline flooding surfaces related to paleolandscape preservation

Within the southeastern Texas coast in the region of the McFaddin Beach Site, Site 41 JF 50, is an 

example of the relatively gentle submergence of Pleistocene coast-parallel terraces by rising sea level. 

Behind the modern shoreline, the terraces are covered as sea level rises by a blanket of marsh that allows 

for little, if any disturbance of the surface, except by bioturbation. Elsewhere, where a narrow lagoon 

exists between a barrier island and landward edge of the lagoon, there often exists a marsh fringe that also 

protects the terrace as it is submerged. Similar conditions can be observed along the inland edge of the 

Louisiana Chenier Plain, where the Pleistocene coast-parallel terraces, along with intact pimple mounds, 

are disappearing beneath coastal marsh.

Within estuaries, the effects of the transgression of the bayline flooding surface can be highly variable,

depending on factors such as relief and steepness of slope of geomorphic surfaces being submerged and 

fetch of the bay. In case of fluvial terraces, bayhead deltas, and floodplains, found within fluvial valleys, 

their generally flat nature allows large areas of them to be rapidly flooded in response to relatively small 

rises in relative sea level (Rodriguez et al. 2005). This would allow for rapid flooding, submergence, and 

subsequent preservation of various geomorphic surfaces and associated cultural deposits within valleys 

during rises in relative sea level. Similarly, the episodic periods of rapid, relative sea level rise noted by 

Milliken et al. (2008b and 2008c), Rodriguez et al. (2005), and others should have the same effect of 

rapidly submerging relatively flat to gently sloped geomorphic surfaces and ultimately abetting their 

preservation with paleovalley fills. Finally, the aggradation of a prism of fluvial sediment upriver from 

bayline should have the effect of burying and preserving the surfaces of crevasse splays, natural levees, 

and geomorphic surfaces well in advance of the flooding of a river valley.

On the other extreme, cultural resource studies, e.g., Lowery (2001) and Lowery and Stanford (2013), of 

large estuaries, such as Chesapeake Bay, Maryland, have found that strong wave action can be generated 

where the over-water distances across which the wind blows or “fetch” is large enough and the orientation 

to prevailing winds is optimum. Such wave action can result in severe erosion of the shoreline and 
43

destruction of cultural resources during coastline flooding. In addition, where wave action is parallel to 

the eroding shoreline, littoral drift has the potential to move artifacts great distances along the shorelines.

In the case of the Mississippi River Delta, the effect of soundline or bayline flooding and associated land 

loss has long been noted by numerous studies. However, except for briefly noting the reduction of shell 

and other middens into erosional beach lags, precise effects of soundline or bayline flooding on coastal 

sites has been noted by only a relative few studies, such as Gagliano (1984), Gagliano et al. (1982), 

Perrault et al. (1994), Weinstein et al. (2012), and Saltus et al. (2003). Although only as little as the upper 

0.5 to 1 meter of the deltaic surface might be eroded, it is enough for waves and currents to winnow out 

finer-grained material from either shell middens or earth mounds, earth middens, and other cultural 

deposits. This process destroys the integrity of a site in terms of stratigraphy and geometry, and reduces it 

to a reworked, erosional lag of the coarsest components, such as shell and artifacts. These components are 

often thrown back on the marshy shore as a perched beach composed of a thin veneer of shell and other 

site debris. Because of subsidence within the delta plain, additional intact cultural deposits often occur on 

the level of shoreline erosion and survive either soundline or bayline flooding (Gagliano 1984, Gagliano 

et al. 1982, Perrault et al. 1994, Saltus et al. 2003). 

Judging from what is currently known about coastline flooding, it quite likely that significant preservation 

of geomorphic surfaces and associated cultural resources occurs during this early stage of a transgression

as the bay-shore processes encroach on paleolandscapes. Even if some disturbance of geomorphic 

surfaces occurs, it appears that a significant proportion of extant cultural resources survive. However, as 

discussed later, the ultimate preservation of geomorphic surfaces and associated cultural resources 

depends how much they are impacted by subsequent shoreface erosion during marine transgression.

5.2.2 Transgression–shoreface ravinement

As pointed out by Pearson et al. (1986) and Waters (1992), the most destructive part of a transgression, in 

terms of geomorphic surfaces and cultural resources, is the result of shoreface erosion and corresponding 

formation of a ravinement surface. As previously noted, shoreface erosion along the gulfward edge of 

barrier islands, chenier plains, and barrier island arcs will result in the removal by erosion of a substantial 

thickness of the uppermost sedimentary deposits and the cultural deposits that they contain forming these 

landforms (Figure. 4.1). Finer-grained material would be removed in suspension and coarser material

would be transported offshore by geostrophic currents, and along the coast by longshore currents, and 

deposited along the shore as a transgressive beach. Some of the denser and larger material, e.g., shells, 

artifacts, and pedogenic nodules, tend to remain in place and concentrate as an erosional lag associated 

with and immediately overlying the ravinement surface. Some of the less dense and smaller artifacts 

would also be thrown onshore and incorporated into transgressive beaches and beach ridges.

Saltus et al. (2003) and Murphy (1990) report the presence of submerged, in situ cultural deposits 

gulfward of the eroding shoreline of coastal barrier islands that are associated with either a submerging 

delta plain or coast-parallel terrace. From these observations, both concluded that cultural deposits can 

survive shoreface erosion. However, both studies are looking at the beachface, which is only in the initial 

phase of erosion, and uppermost segment of the entire shoreface. Because shoreface erosion is as much as 

6 m to 15 m along the Texas coast (Rodriguez et al. 2001, Wallace et al. 2010); 8 m to 12 m along the 

front of the Chandeleur Islands (Miner et al. 2009), and 3 m to 5 m along the Isles Dernieres (Penland et 

al. 1985), it appears that shoreface erosion effectively destroys most geomorphic surfaces and associated 

cultural deposits. The main preservation of paleosurfaces occurs in paleotopographic features, e.g.,

paleovalleys and in the northeastern Gulf of Mexico offshore Florida, sinkholes, and paleokarst valleys, 

that lie below the level of shoreface erosion (Figure 5.4) (Pearson et al. 1986, Faught 2004).
44

The McFaddin Beach Site (41JF50) is a well-documented example of a mixed assemblage containing 

Clovis, Eden, San Patrice, and other projectile points, along with other artifacts and Late Pleistocene 

mammoth, camel, and horse bones concentrated on a transgressive beach by shoreface erosion of 

Holocene and Late Pleistocene deposits (Coastal Environments, Inc. 1977, Long, 1977, Stright et al. 

1990).

Figure 5.4. Preservation potential of cultural deposits within a major paleovalley fill (e.g., Sabine) 

in respect to depth of erosion of the ravinement surface. 

Modified from Pearson et al. (1986).

Although the bulk of the barrier island and lagoonal sedimentary sequences underlying the barrier island 

and associated lagoon are preserved from shoreface erosion, the beach ridges and other landforms

underlying surficial deposits are eroded and reworked. It is highly unlikely that any cultural deposits 

associated with beach ridges and other landforms can survive a shoreface retreat intact.

5.2.3 Tidal ravinement

The formation of a tidal ravinement also has the potential to disturb and destroy paleosurfaces and 

associated cultural deposits. Anchored tidal ravinements, which are associated with the inlets of barrier 

arcs of subsiding sublobes of the Mississippi River Delta, widen and deepen the paleochannels of their 

former deltaic distributaries. As previously noted, the result of overall enlargement of such paleochannels 

is to erode and disperse with time the natural levees of and any cultural deposits associated with the 

paleochannels. Similarly, the creation of a sweeping tidal ravinement by the lateral migration of a tidal 

inlet will erode the sequence of sediments that lie above the depth of the thalweg of the tidal channel as it 

migrates. This depth of tidal ravinement can be greater than that of shoreface ravinement at deep inlets, if 

the inlet is deeper than the depth of shoreface erosion (Miner et al. 2007). At the least, any beach ridges 

and typically submerged or buried natural levees of a deltaic distributary paleochannel will be destroyed,

along with any associated cultural deposits as a tidal inlet migrates through them and a laterally extensive 

tidal ravinement surface is left behind. At worst, the tidal channel erodes deeply enough to cut entirely 

through the coastline flooding surface and into either channel belt sands of a paleovalley fill or sediments 

from a previous sea level cycle.
45

As reported by Saltus et al. (2003), an example of the ongoing destruction of a paleosurface and 

associated cultural resources is the Fort Livingston Site (16JE49) that was constructed at the mouth of 

Barataria Bay and the western end of Grand Terre Island. On the westernmost end of this island, 

numerous ancillary buildings and a rail line were built on a sand dune spit during the construction of this 

fort. By the 1930s, enlargement of the tidal inlet had completely eroded this spit and converted this area to 

open water that later refilled with sand as the spit was rebuilt. Magnetometer investigations by Saltus et 

al. (2003) in this spit and adjacent tidal channel found magnetic expressions that are likely associated with 

the heavier remnants of these historic structures. These cultural materials were likely dropped down by 

erosion to the bottom of the tidal channel when these structures were destroyed by enlargement of the 

tidal channel and later reburied when the spit reformed. McNinch et al. (2006) and Quinn (2006) describe 

the scour and settling processes that would cause the downward movement of individual artifacts by 

removal of sand and smaller-sized sediment. 

5.2.4 Normal regression–progradation

As discussed by Waters (1992), normal regression, as the result of progradation of strand plains, chenier 

plains, and barrier islands is most conducive to the creation and preservation of geomorphic surfaces and 

associated cultural resources. As the shoreline builds gulfward, their associated landforms are temporarily 

removed from destructive coastal processes as former coastal sites are left further and further inland as the 

shoreline progrades gulfward. If progradation is periodically interrupted by a cessation in sediment 

supply, they will be subjected to wave erosion and shoreline retreat and shoreface erosion. This process 

can produce transgressive beach ridges and unconformities in beach ridge sequences. The transgressive 

beach ridges can contain a mixture of older cultural material that has been reworked from the older 

deposits (Coastal Environments, Inc. 1977, Gagliano 1984, Gagliano et al. 1982, Waters 1992).

These prograding shorelines create an ordered sequence of beach ridges that become progressively 

younger toward the shoreline on which lie archaeological sites that overall also become progressively 

younger toward the shoreline (Figure. 5.2.3). The basic cause of this chronological ordering of 

archaeological sites is that prehistoric people found it advantageous to locate their sites near the ocean. 

Such chronological ordering of archaeological sites has been documented in the chenier plain of western 

Louisiana. Within it, the distribution of initial occupation sites corresponds to the sequential development 

of major chenier ridges (McIntire 1971, Coastal Environments, Inc. 1977, Gagliano 1984, Gagliano et al. 

1982, Waters 1992). However, as previously noted, regardless of whether a set of beach ridges are 

gradually or rapidly submerged, it is quite unlikely that they and associated cultural resources will survive 

intact. 
46

Figure 5.5. Distribution of initial occupation sites in a prograding beach sequence. From A-oldest; 

to E-youngest. 

Reprinted from Gagliano et al. (1982).

Similarly, the Holocene progradation of the Mississippi River Delta created an environment conducive to 

the creation and preservation of geomorphic surfaces and associated cultural resources (McIntire 1971,

Coastal Environments, Inc. 1977, Gagliano et al. 1982, Waters 1992). After establishment of a prodelta 

platform, deltaic distributaries prograde and bifurcate, which results in the active construction of the plain 

of a delta lobe. At this time, permanent settlements are lacking from this delta plain because it is too 

prone to flooding to be comfortably settled. The active construction of a delta plain continues until the 

distributary course is no longer hydraulically efficient. Then it is abandoned for a new trunk channel that 

starts construction of a new delta lobe. After a delta lobe is abandoned, , its plain is actively occupied 

because of the reduced flooding and greater productivity and diversity in fauna and flora. In addition, the 

abandonment of a delta lobe initiates the transgressive phase of the delta cycle. Initially, as the delta plain 

subsides, fine-grained sediments accumulate and slowly bury natural levees and mouth bar accretionary 

ridges and their associated cultural deposits in a protective mantle (Figure 5.6). An example of such a site 

is the Bayou Jasmine Site (16SJB2), which consists of Rangia shell and earth middens, is situated on 

natural levee ridges of Mississippi River distributaries, and is now subsided as much as 5.5 m and lies 

buried beneath fine-grained deltaic sediments. However, as the abandoned delta lobe subsides, coastal 
47

processes rework the gulfward margin of the delta plains and create a sandy barrier shoreline backed by 

bays and sounds. At first, the surficial deposits and cultural resources are disturbed and reworked by 

coastline flooding (Figure 5.7). Later, as the barrier islands move inland, the deltaic and associated 

cultural deposits are deeply eroded by shoreface processes, as previously described (Figure 4.1). 

Figure 5.6. Shell mound on a subsided natural levee ridge enveloped by marsh mud. 

Reprinted from Gagliano et al. (1982).
48

Figure 5.7. Schematic dip cross section, central coastal plain, Texas.

Reprinted from Saltus et al. (2003).
49

5.2.5 Forced regression

Forced regressions resulting from the fall of relative sea level would have been very conducive for the 

creation and preservation of geomorphic surfaces and associated cultural resources (Figure 4.4). 

However, they would have been created during the last glacial cycle at times before prehistoric Native 

American occupation of the northwestern Gulf of Mexico Coastal Plain. Although the relict landforms 

created by forced regressions might have later been ideal locations for Native American occupations, 

many of the relict landforms created by forced regression would have largely occupied the relatively flat, 

coast-parallel terraces between major river valleys. As a result, they would be destroyed along with 

associated cultural resources by shoreface erosion when relative sea level rose. Only the landforms 

associated with strips of coastal plain preserved from shoreface erosion by back-stepping of barrier 

islands and fluvial terraces within paleovalleys might have survived relatively intact and undisturbed.
50

6. Stratigraphy of Study Area

As discussed by Young et al. (2010 2012), considerable effort has been made over the last hundred years 

in understanding the stratigraphy and sedimentology of the sedimentary strata underlying the 

northwestern coastal plain of the Gulf of Mexico. The earliest geologic studies, e.g., Deussen (1914 and 

1924), Barton (1930), Trowbridge (1932), Plummer (1932), Price (1934), Weeks (1933 and 1945), and 

Fisk (1940 and1948), focused on outcrop description and correlation. Later studies, e.g., Doering (1935

and 1956), Bernard (1950), Holland et al. (1952), Fisk and McFarlan (1955), Bernard and LeBlanc 

(1965), and Otvos and Howat (1997), incorporated the concept by Fisk (1940 and 1944) of a direct oneto-one correlation between sea level fluctuations, coastal-parallel terrace formation, and only four major 

continental glaciations, the Nebraskan, Kansan, Illinoian, and Wisconsinan stages of Kay and Apfel 

(1929) and Reed and Dreeszen (1965). The combination of these concepts resulted in the arbitrary and 

ultimately temporary subdivision of the Lissie Terrace into Bentley and Montgomery terraces within 

southwest Louisiana and southeast Texas (Heinrich 2009). Some studies, e.g., Akers and Holck (1957) 

and McFarlan and LeRoy (1998a and 1998b), attempted to incorporate subsurface data, primarily 

geophysical logs of oil and gas well and proprietary lithologic and paleontologic data, into their mapping 

and subsurface cross-sections that extended offshore. Although discredited by Boellstorff (1973 and 

1978), these studies based their correlation on the incorrect presumption that there had been only four 

major continental glaciations during the Pleistocene. Most recently, only a set of three coast-parallel 

terraces has been consistently recognized and mapped along the northwestern Gulf of Mexico Coastal 

Plain (Winker 1979 and 1990, Snead and McCulloh 1984, DuBar et al. 1991, Heinrich 2009).

In Texas, this research was evaluated and is summarized by the Bureau of Economic Geology in its 

Geologic Atlas of Texas series (Aronow 1975, Shelby et al. 1968, Proctor et al. 1974, Aronow and Barnes 

1975, Aronow et al. 1975, Brewton et al. 1976a, Brewton et al. 1976b). This mapping was revised in their 

Environmental Geologic Atlas of the Texas Coastal Zone series (Brown et al. 1977 and 1980, Fisher et al. 

1972, McGowen et al. 1976a and 1976b. 

In Louisiana, this research was evaluated and summarized in large-scale geologic maps. First, Snead and 

McCulloh (1984) remapped units and revised stratigraphic nomenclature for their 1:500,000 scale state 

geologic map. In their mapping, terraces were inferred to represent unconformity-bounded “complexes”

composed of sedimentary strata. In addition, the Bentley and Montgomery terraces were collapsed into a 

single Intermediate complex in the absence of recognizable distinctions in terrace surfaces, terrace scarps, 

bounding unconformities, and other features. Later, Snead et al. (1999) and McCulloh et al. (2003) 

reclassified the informal complexes nomenclature into more formal allostratigraphic units. Their 

classification formed the basis for the nomenclature used in the compilation of 1:100,000 geological maps 

(Heinrich et al. 2002 and 2003, Heinrich 2005b and 2006).

The overall conclusion of these studies is that multiple unconformity-bounded, seaward dipping, clastic 

wedges, ranging in age from Pliocene to Late Pleistocene, underlie the coast-parallel terraces that 

compose the northwest Gulf of Mexico Coastal Plain (Figure 6.1). These clastic wedges consist of 

nonmarine sediments that grade gulfward into coastal and eventually shallow marine sediments. In updip 

areas, each of these clastic wedges erosionally truncates and onlaps an underlying clastic wedge. Thin 

erosional remnants, isolated terraces, onlapping veneers, and Holocene alluvial cover make it difficult to 

establish regional correlations between outcropping and subsurface stratigraphic intervals (Winker 1979,

Winker 1990, DuBar et al. 1991).
51

Figure 6.1. Regional dip cross section showing relationships between outcropping formations and 

subsurface stratigraphy, northwestern Gulf of Mexico Coastal Plain. 

Reprinted from Young et al. (2012).

6.1 Allostratigraphic Approach to Pleistocene Unit Definitions

Traditional sedimentary units in geology are defined by their stratigraphic position and their lithology. 

This is effective in older hard-rock geology with extensive beds of consistent rock types. However, the 

largely unconsolidated, heterogeneous, and nondiagnostic nature of the siliciclastic fluvial, deltaic, 

coastal, and shallow marine sediments of the northwestern Gulf of Mexico Coastal Plain that were 

deposited within dynamic and shifting fluvial, deltaic, and coastal environments created widely varying 

packages of sediment juxtaposed among each other. All of these deposits consist of complex 

interfingering depositional facies that are composed of varying proportions of sand, silt, clay, peat, and 

minor amounts of gravel (McCulloh et al. 2003, Heinrich et al. 2015).

As a result, traditional lithologic formation categories are of less value here and Quaternary geologists 

must use other distinguishing criteria. In the subsurface, it becomes useful to redefine them as 

alloformations, where the package of sediment is defined by its bounding unconformities. An allogroup is 

several related alloformations. Also, sometimes it is useful to subdivide these alloformations into 

allomembers. Where exposed at the surface, the upper bounding surface of an alloformation typically 

forms a distinctive coast-parallel or fluvial terrace surface that can be recognized and mapped using 

topographic and pedologic data (Snead et al. 1999, Young et al. 2012, Heinrich et al. 2015).

By the late 1980s, the Louisiana Geological Survey, e.g., Autin (1989), started examining the application 

of allostratigraphic concepts and nomenclature to the definition and mapping of surface Plio–Pleistocene 

units. It was eventually concluded that allostratigraphy offers an effective, if not essential, approach to 

their delineation and classification. It was also noted that Plio–Pleistocene units of northern Gulf of 

Mexico Basin subsidence exhibit a clear spectrum of preservation of surface morphology varying from 

pristine relict and active landforms of the Holocene coastal plain to deeply dissected terrace surfaces of 

the oldest Pleistocene allostratigraphic units. In case of older Pliocene strata, the original constructional 

surfaces have been destroyed by erosion. As a result, allostratigraphic nomenclature has been heavily 

used in the STATEMAP-funded geologic mapping projects of the past two decades to map approximately 

three-fourths of the surface of Louisiana that is covered by Quaternary strata (McCulloh et al. 2003).
52

6.1.2 Beaumont Alloformation (Prairie Allogroup)

The Prairie Allogroup, formerly called the “Prairie complex” or “Prairie Formation,” is a collection of 

Pleistocene depositional sequences of alloformation rank (Autin et al. 1991, McCulloh et al. 2003,

Heinrich et al. 2002, Heinrich et al. 2003, Heinrich 2005b, Heinrich 2006). The sediments of the Prairie 

Allogroup accumulated within a diverse variety of coastal plain settings, i.e., fluvial (meander-belt, 

backswamp, and braided-stream), colluvial, estuarine, deltaic, and shallow marine environments. It had 

been argued that these sediments largely accumulated over a considerable part of the late Pleistocene 

(Autin et al. 1991, Otvos and Howat 1997, Otvos 2005, McCulloh et al. 2003, Heinrich 2009). More 

recently, correlation of the unconformity at the base of the Beaumont Alloformation by Young et al. 

(2012) indicates that it accumulated during a large part of the middle Pleistocene and the earliest part of 

the late Pleistocene.

The surface of the Prairie Allogroup is a coast-parallel terrace that extends along the northwest coast of 

the Gulf of Mexico from a point about 110 km (67 mi) south of the Rio Grande in Mexico over to at least 

Mobile Bay, Alabama. This coast-parallel terrace is the lowest continuous terrace and lies directly above 

Holocene coastal and flood plains (Figures 3.3 and 6.2.1.1). Unlike older coast-parallel terraces, the 

surface of the Prairie Allogroup exhibits recognizable constructional topography. However, the relict 

landforms that it exhibits are distinctly more poorly preserved than those seen on the fluvial terraces of 

the Deweyville. It comprises multiple stratigraphic units of alloformation rank (McCulloh et al. 2003,

Heinrich et al. 2002 and 2003, Heinrich 2005b, Heinrich 2006).

The Beaumont Alloformation is an unconformity bounded stratigraphic unit. It is separated from the 

underlying Lissie Alloformation by a regional unconformity that has been defined and mapped by Young 

et al. (2010 and 2012). Except where cut by valleys formed during sea level lowstands of the last glacial 

cycle, the upper boundary of the Beaumont Alloformation consists of the surface of the coastal plain.

It was originally named the “Beaumont Clay” by Hayes and Kennedy (1903) for the City of Beaumont, 

Jefferson County, Texas. There is a lack of any type section because the original description of the 

Beaumont Clay is a composite compiled from subsurface information from numerous wells. 

Since it was named by Hayes and Kennedy (1903), the Beaumont Alloformation was intensively studied 

as either the Beaumont Clay or Beaumont Formation in the southeast Texas coastal plain by various 

geologists (e.g., Deussen 1914, Deussen1924, Barton 1930, Aronow 1971, Aronow 1986, Aronow and 

Barnes 1975, Fisher et al. 1972, Shelby et al. 1968, Van Siclen 1985). It was first mapped across most of 

the Texas Coastal Plain by Darton et al. (1937) and later in Southwestern Louisiana as the either the 

Prairie Formation or Prairie Terrace(s) (e.g., Fisk 1944, Bernard 1950). Doering (1956) subdivided it into 

the Eunice and Oberlin formations. More recently, Snead and McCulloh (1984), Mossa and Autin (1989), 

and Autin et al. (1991) recognized it as an informal allostratigraphic unit called the “Prairie complex.” In 

recognition of the persistence of the name “Beaumont” for more than 80 years and rules of priority, Snead 

et al. (1999) and McCulloh et al. (2003) redesignated the “Prairie complex” as the Beaumont 

Alloformation in southwestern Louisiana. This is the nomenclature that has been used in Louisiana 

Geological Survey geologic mapping. Most recently, the Texas Water Development Board, e.g., Young et 

al. (2010 and 2012), redefined it as an unconformity bounded stratigraphic unit. 

As indicated by its original name, the Beaumont Alloformation is a predominately fine-grained 

stratigraphic unit that consists largely of varicolored, laminated to massive, calcareous silty clays that 

often contain calcareous nodules (Figure 6.2). The clays contain organic matter, partly decomposed tree 

remains (predominantly cypress), and shell beds (Rangia cuneata and Crassostrea spp.). Other animal 

fossils include mammoths, horses, and an indigenous foraminiferal fauna. These clays enclose brown, 

yellow, and blue sands and silty sand that occur as dip-oriented, narrow, ribbon-like channel belts and 
53

deltaic-distributary channels that are 6 to 15 m thick and are locally stacked as thick as 45 m. Sandy 

sediments also comprise a coast parallel sandy beach ridge complex known as the Ingleside barrier and/or 

strandplain system. The interbedded muddy units are generally of similar thickness to the sands (Hayes 

and Kennedy 1903, Price 1934, Knox et al. 2006, Young et al. 2012).

Figure 6.2. Regional dip cross section showing relationships between stratigraphic units within 

the study area and onshore Southwest Louisiana.

Except where cut by river valleys that formed during sea level lowstands of MIS 4, 3, and 2 and the 

Lafayette meander belt, the upper contact of the Beaumont Alloformation is known informally as the 

“Beaumont terrace” or “Beaumont surface” (Figure 3.2). This surface is an extensive coast-wise terrace 

that extends from the western valley wall of the Mississippi River Alluvial Valley past the Rio Grande to 

the Tamaulipas Range in northeastern Mexico (Bernard 1950, Doering 1956, Winker 1990). The surface 

of the Beaumont Alloformation exhibits moderately well preserved relict depositional topography of 

regional rivers and small streams (Fisk 1948, Fisk and McFarlan 1955, Aronow 1986, Winker 1990,

Saucier and Snead 1989). It also includes extensively modified, coast-parallel, sandy beach ridges often 

known as the Ingleside Shoreline, Ingleside barrier system, or Ingleside strandplain system. Within 

Southwest Louisiana, the crest of this ridge ranges in elevation from 9 m to 11 m above mean sea level 

and rises 2 m to 4.5 m above the adjacent surface of the Beaumont terrace (Heinrich 2007). The 

Beaumont terrace is generally flat, little-dissected, and slopes gently. Within Southwest Louisiana, its 

slope is about 0.2 m to 0.3 m / km and 20 km to 28 km wide in the Lake Charles area. This surface and 

associated relict landforms have been extensively modified by pedogenesis and pedoturbation, resulting 

in its upper part being significantly weathered and overconsolidated (Fisk 1948, Fisk and McFarlan 1955,

Johnson et al. 2008). The alteration of the surface of the Beaumont Alloformation has also resulted in the 

formation of a variable combination of innumerable pimple mounds and “blow-out” depressions along 

with extensive gilgai topography (Plummer 1932, Johnson et al. 2008). This surface extends beneath 

terminal Pleistocene to Holocene coastal-plain deposits and marine deposits and extends offshore beneath 

the Texas and Louisiana continental shelves. Winker (1990) and DuBar et al. (1989) correlate it with his 

R1 reflector. Within interfluves and uplands between Holocene and Wisconsinan paleovalleys and 

channel belts, it forms part of the Holocene-Pleistocene surface. 
54

Until recently, there has been a lack of any agreed-upon definition of the base of the Beaumont 

Alloformation, as noted by Aronow (1971). Winker (1979) has projected its position to a seismic reflector 

offshore that he regards as correlative to the top of the Lissie Formation. Although this inferred base may 

prove to be valid, the documentation does not constitute a satisfactory definition. More recently, Young et 

al. (2010 and 2012) define the lower contact of the Beaumont Alloformation as a regionally extensive and 

laterally extensive flooding surface that is correlated with the Trimosina A micropaleontological zone 

offshore at about 0.6 Ma (Figure 6.3). Updip, this flooding surface is correlated northward along the bases 

of major channel sands to where it crops out and the Beaumont Alloformation onlaps on the surface of the

Lissie Alloformation (Young et al. 2012). 
55

Figure 6.3. Correlation chart of regional Late Neogene strata showing the relationship between biostratigraphy, coastal onlap curve, and 

marine oxygen isotope stages. 

lcha = Lava Creek Ash and Qts = Terrebonne Shale. Because of scale, Deweyville Allogroup (green) and Holocene units are shown in Figure 6.4.
56

Contrary to the assumptions of previous studies, the Beaumont Alloformation consists of sediments that 

accumulated during multiple high-frequency glacio-eustatic sea-level fluctuations (Figure 6. 4). Onshore, 

the bulk of its deposition occurred during interglacial highstands of sea level in fluvial, deltaic, and 

marginal-marine systems (Blum et al. 2013). Within Texas, the highstand positions of the Beaumont 

shoreline almost coincide with those of the modern shoreline in Texas and as far as 60 km inland in 

Southwest Louisiana (Solis 1981, Knox et al. 2006, Heinrich 2007, Young et al. 2012). 
57

Figure 6.4. Late Pleistocene stratigraphy of Southwest Louisiana and offshore study region 

showing the relationship between major coastal plain stratigraphic units, global sea 

level, paleomagnetic chronology, and chronostratigraphy.

The Beaumont Alloformation has been long considered to be late Pleistocene in age (DuBar et al. 1991). 

Pleistocene-age fossils have been found in the Beaumont Alloformation at numerous locations on the 
58

Texas Coastal Plain (Maury 1920, Maury1922, Plummer 1932, Price 1934). Initially, geologists H.N. 

Fisk and his associates (e.g., Fisk 1938, Fisk 1944, Fisk and McFarlan 1955, Bernard and LeBlanc 1965) 

assigned a mid-Wisconsinan age to the Beaumont Alloformation that corresponds approximately to MIS 

3. Later, Mange and Otvos (2005) and Otvos (2005) obtained optically-stimulated luminescence ages 

from the Beaumont Alloformation that ranged between Eowisconsin (MIS 5d–5a) and Wisconsinan 

(MISs 4 and 3) ages. However, optically-stimulated luminescence ages from the Beaumont Alloformation 

and correlative Hammond Alloformation indicate that the Beaumont Alloformation is older than 

Wisconsinan (MIS 4 and 3) and its surface sediments accumulated during marine isotope stages 5e and 

5a. Optically-stimulated luminescence ages from younger Deweyville Allogroup and Mississippi River 

valley train deposits substantiated this interpretation (Shen et al. 2012 and 2013).

Furthermore, Simms et al. (2013) obtained optically-stimulated luminescence ages of 119 to 127 ka from 

samples from depths between 1.5 m and 4.0 m beneath the surface of the Ingleside strandplain. Samples 

between depths of 1.0 m and 1.5 m beneath the surface of the Ingleside strandplain yielded ages of 1.32 to 

57 ka. These ages were interpreted as representing later eolian reworking of the surface of the Ingleside 

strandplain. 

Within Southwest Louisiana, the position of relict channel belts gulfward of the Ingleside Shoreline and 

the presence of incised fluvial ridges and beheaded deltaic distributary ridges as seen in LiDAR DEMs, 

topographic maps, and aerial photographs indicates that the final accumulation of sediments comprising 

the Beaumont Alloformation occurred during the falling stages of marine isotope stage 5. The Beaumont 

Alloformation clearly predates the terraces of the Deweyville Allogroup that lie within the valleys of the 

Sabine, Trinity, Calcasieu, and other rivers and the adjacent Lafayette meander belt of the Avoyelles 

alloformation. Downdip correlations by Young et al. (2012) argue that the Beaumont Alloformation 

began accumulating 600,000 years ago.

Within the Beaumont Alloformation, its significant cultural and sand resources will be restricted to its 

upper contact. 

Braid belts and paleovalleys associated with ancient paleochannels of either the Red, Calcasieu, 

Mississippi, and other rivers contain large quantities of sand. However, only those fluvial sand bodies at 

the upper contact of the Beaumont Alloformation and underlying a thin blanket of marine sediments 

would be shallow enough to be economic to use. 

Similarly, because of the age of the Beaumont Alloformation, its accumulation predates the accepted 

range of time during which this part of North America was first populated (Stright 1990, Anderson et al. 

2013). As a result, cultural materials can be expected to be restricted to the the surface of this stratigraphic 

unit and within paleosols, and have been buried by younger deposits, e.g., cover sands. Preferred 

locations for the occurrence of archaeological sites on the surface of the Beaumont Alloformation include 

the tops of pimple mounds and within the sandy soils associated with them (Aten and Bollich 1981; Jones 

and Shuman 1988), the relict natural levees of Pleistocene paleochannels, and terrace edges overlooking 

valleys cut into the Beaumont surface by modern streams and bayous (Weinstein et al. 1979b, Coastal

Environments, Inc. 1977).

6.1.3 Avoyelles Alloformation

The Avoyelles alloformation is an informal unconformity-bounded stratigraphic unit. At its base, it is 

separated from the underlying Beaumont Alloformation by an extensive unconformity that represents the 

basal scour surface of a MIS 3 Mississippi River meander belt. The upper contact of the Avoyelles 

alloformation is a surface that includes the channel belt of the Lafayette Meander Belt and associated 

flood basin (Figure 3.2). This surface is covered by a westward thinning blanket of Late Wisconsinan 

(MIS 2) Peoria Loess that ranges from about 7 m in thickness at the western valley wall of the Mississippi 
59

Alluvial Valley to about 1 m in thickness within a distance of about 8 km (Miller et al. 1985, Saxton 

1986, Callihan 1988, Rouly 1989, Mateo 2005).

The Avoyelles alloformation was informally named by by Snead et al. (1999) and McCulloh et al. (2003) 

for Prairie des Avoyelles and Avoyelles Hills within Avoyelles Parish, Louisiana. These are the 

northernmost occurrence of the alloformation in Louisiana and where Fisk (1940) described and 

recognized the significance of the surface morphology of the Avoyelles alloformation.

Currently, there is no formal type section for the Avoyelles alloformation. However, two cores that 

partially penetrated it within the Avoyelles Prairie are described in detail by Autin and Aslan (2001). 

Also, Mateo (2005) illustrates four cores, nos. 0394-006, 0494-022, 0494-023, and 0494-024, from a 

point bar of the Lafayette Meander Belt near Carencro, Louisiana that penetrated the Avoyelles 

alloformation. Lithologic logs of geotechnical borings that penetrate the Avoyelles can be found in 

Appendix 4 of Rouly (1989) and in the Type III Solid Waste Permit Application for the ANGCO facility 

near Carenco, Louisiana (KourCo 2011).

The most detailed information about the lithology of Avoyelles alloformation comes from the borings of 

Mateo (2005) taken from a point bar of the Lafayette Meander Belt north of Opelousas and geotechnical 

logs of the ANGCO facility near Carenco, Louisiana. They show that beneath the Peoria loess, the 

channel belt deposits of the Avoyelles alloformation consist of sand that fines upward and contains 

occasional ripple and parallel laminations and interbeds of silty loam. The uppermost unit of the 

Avoyelles alloformation consists of a 1.5 m to 6 m-thick, massive, gray to brown silty clay loam, which 

occasionally includes a weakly developed paleosol. It is well indurated and contains iron and manganese 

mottling. This unit might correlate with the “Yellow Clay” unit of Autin and Aslan (2001). So far, no 

fossils have been reported from the Avoyelles alloformation. 

The surface of the Avoyelles alloformation is a loess-mantled fluvial terrace that forms the southernmost 

segment of the western valley wall of the Mississippi Alluvial Valley (Figure 3.2). This terrace occupies a 

triangular 0 to 22 mi (0 to 35 km) wide area, which extends eastward from the western valley wall of the 

Mississippi Alluvial Valley from New Iberia, Louisiana to Abbeville, Louisiana, and just west of 

Intracoastal City, Louisiana. This fluvial terrace consists of a Pleistocene Mississippi River channel belt 

that truncates older, loess-covered fluvial-deltaic plains of the Red River, which are underlain by the 

Beaumont Alloformation (Saucier 1994, Mateo 2005, Shen et al. 2012). Its slope is difficult to determine 

because of faulting and the variable thickness of its loess cover. However, the analysis of available 

LiDAR DEMs where the loess is thinnest suggests that it slopes gulfward at about 0.07 m per kilometer. 

Just before it disappears beneath the Holocene sediments of the Mermentau Alloformation, this channel 

belt bifurcates into distinct channel belts. Both channel belts exhibit moderately to poorly preserved relict 

constructional landforms, relict river courses, meander loops, and ridge and swale topography that are 

comparable in size to those of the modern Mississippi River (Fisk 1948, Jones et al. 1954, Saucier 1994,

Shen et al. 2012). South and southeast of Jefferson Island and within Iberia Parish, a featureless and flat 

clay plain, a relict flood basin, forms the surface of the Avoyelles alloformation between the Lafayette 

Meander Belt and the adjacent Holocene delta and Chenier Plains (Autin 1984). In places, the surface of 

this relict meander belt and overlying loess blanket have been displaced by faulting to form east-west 

trending fault-line scarps (Heinrich 2006, Heinrich et al. 2003). 

The base of this allostratigraphic unit is likely a composite and diachronous erosion surface created by the 

slight incision and lateral migration of an ancient Mississippi River. The sediments underlying it are Red 

River deposits of the Beaumont Alloformation.
60

The Avoyelles alloformation consists largely of thick point bar, crevasse, overbank, and other fluvial 

sediments. The channel deposits are overlain by 2 m to4 m of Mississippi River overbank deposits. In 

turn, the overbank sediments are veneered by 2 m to5 m of Peoria Loess (Saucier 1994, Mateo 2005, 

Shen et al. 2012). The sediments that comprise the Avoyelles alloformation accumulated within the Late 

Pleistocene Lafayette meander belt and associated flood basin (Heinrich 2005b). The ancient delta 

associated with this meander belt now lies underwater offshore of Cameron Parish where it forms part of 

the southern edge of the Louisiana Continental Shelf. Likely, one of the meander belts terminates with the 

West Louisiana Delta of Wellner et al. (2004).

The age of the Avoyelles alloformation is constrained by stratigraphy, optically stimulated luminescence 

ages, and thermoluminescence ages. it is incised into and younger than Red River deposits of the 

Beaumont Alloformation that are best interpreted to date to MIS 5a (Shen et al. 2012, Shen et al. 2013). 

They are overlain by MIS 2 loess of the Peoria Loess (Bettis et al. 2003, Mateo 2005). Excluding a 25 ± 3 

ka age that is anomalous relative to the age of the overlying Peoria Loess, using optically stimulated 

luminescence, Mateo (2005) dated the Avoyelles alloformation at about 55 ± 6 ka. This age is consistent 

with the cross-cutting relationship between the Beaumont and Avoyelles alloformations and ages obtained 

from the Marksville segment of the Avoyelles alloformation. Optically-stimulated luminescence ages of 

85 ± 10 and 86 ± 14 ka and a thermoluminescence age of 54 ± 6 ka were reported by Otvos (2005) from 

borrow pits in the northern and southwestern sector of the Lafayette Meander Belt. However, it is hard to 

interpret them in light of the stacking of the deltaic, meandering river and loess deposits in this area 

because the stratigraphic context for the dated samples is lacking.

Offshore, the Avoyelles alloformation would be a significant source of high quality sand for coastal 

restoration projects. As indicated by the core descriptions of Mateo (2005) and KourCo (2011) the 

channel belt sand deposits of it consist of a thick layer of well-sorted (poorly-graded) sand that because of 

shoreface erosion should lie relatively close to the surface of the gulf bottom. 

As in case of the Beaumont Alloformation, the deposition of Avoyelles alloformation predates the 

accepted range of time during which this part of North America was first populated (Stright 1990,

Anderson et al. 2016). As a result, cultural materials can be expected to be restricted to the surface of this 

stratigraphic unit and within paleosols, and to have been buried by younger deposits, e.g., possibly loess. 

Preferred locations for the occurrence of archaeological sites would be the the relict natural levees of 

Pleistocene paleochannels, and terrace edges overlooking valleys cut into the Lafayette Meander Belt by 

modern streams and bayous (Weinstein et al. 1979a, Weinstein 1981, Coastal Environments, Inc. 1977).

6.2 Trinity-Sabine Unconformity

Both onshore and in the study area, the Deweyville Allogroup is separated from the older deposits of the 

Prairie Allogroup that form the base and sides of paleovalleys cut into these older deposits (Figure 6.5). 

The base of the valley is a diachronous, composite surface formed by at least three cut-and-fill events and 

subsequent lateral migration of the associated fluvial system (Sylvia and Galloway 2006). Each of these 

events deposited a recognizable allostratigraphic package of fluvial sediments bounded on the sides and 

base by erosional surfaces and on their top by a terrace surface (Blum and Valastro 1994, Blum et al. 

1995). Within the study area, the basal unconformity of these paleovalleys is known informally as the 

“Trinity-Sabine unconformity.”

Contrary to the reconstructions of the prehistory of the Mississippi River Alluvial Valley, this surface was 

never subaerially exposed. Thus, it was never available for habitation regardless of the age of formation 

of a specific surface.
61

Figure 6.5. Cross-section across the Trinity-Sabine Paleovalley. 

Modified and reinterpreted from Thomas (1991) showing major bounding unconformities and 

allostratigraphic units.

6.3 Deweyville Allogroup

The Deweyville Allogroup is an unconformity-bounded sedimentary package that underlies subregionally extensive fluvial terraces that lie topographically below the position of the surface of the 

Beaumont Allogroup, the Prairie terrace or Prairie surface, and geomorphically above the modern flood 

plain along streams and rivers of the northern Gulf coastal plain (Bernard 1950) (Figure 6.6).
62

Figure 6.6. Holocene–Pleistocene Surface, 5-m Contour Interval. 

The deep trenches are the paleovalleys of the Trinity (Tr), Sabine (Sa), and Calcasieu (Ca) rivers. Figure created using Global mapper from data 

supplied by Milliken (2010).
63

Bernard (1950) focused on the uppermost of three fluvial terraces in the area of Deweyville, Texas and 

did not designate a type section. Gagliano and Thom (1967) later identified fluvial terraces along much of 

the Gulf and Atlantic coastal plains as the Deweyville terrace. However, terraces and associated fluvial 

sediments compatible with the original definition of Bernard, (1950) are best restricted to streams of 

intermediate size in Texas, Louisiana, Arkansas, and Mississippi, e.g., Fisher et al. (1972), Kesel (1980),

Snead and McCulloh (1984), Winker (1990), Blum et al. (1995), Garvin (2008), and Blum et al. (2013). 

Pearson et al. (1986) recognized and mapped the Deweyville terrace for a distance of up to 48 km south 

of the current shoreline along the offshore Sabine River paleovalley beneath the submerged LouisianaTexas continental shelf. Thomas (1991) also recognized the presence of three buried terraces and 

associated fluvial sediments within offshore Sabine and Trinity River paleovalleys.

The base of this allogroup is a composite and diachronous erosion surface created by the periodic 

incision, lateral migration, and channel belt formation of streams of intermediate size in Texas, Louisiana, 

Arkansas, and Mississippi (Figures 3.3 and 6.5). The scoured base of each of the, typically three, 

Deweyville channel belts progressively truncate the scoured base of an older channel belt to form the base 

of a paleovalley that consists of a composite of erosional surfaces that are younger with depth. The 

exposed surface of the Deweyville typically exhibits well-preserved oversized fluvial features (Sylvia and 

Galloway 2006, Garvin 2008, and Blum et al. 2013).

The top of the Deweyville Allogroup is a flight of typically three fluvial terraces that lie topographically 

below the position of the surface of the Beaumont Allogroup and geomorphically above the modern flood 

plain along streams and rivers of the northern Gulf coastal plain (Figures 3.3). Gulfward along these 

streams, these terrace surfaces are progressively buried by Holocene flood plain deposits that, in turn, are 

buried within paleovalleys by estuarine sediments adjacent to the coastline and shallow marine sediments 

offshore (Gould and McFarlan 1959, Byrne et al. 1959, Milliken et al. 2008b, Milliken et al. 2008c). In 

southeast and central Texas, the terraces of the Deweyville Allogroup are sometimes blanketed with a 

few-meter thick layer of windblown sand, known as the “Holocene sandy mantle,” that has accumulated 

over the Early to Late Holocene. Although thin, these windblown sand deposits have been found to 

contain in situ and stratified archaeological deposits. Where covered by these younger deposits, either 

overbank or eolian, paleosols are commonly found developed within the Deweyville Allogroup (Abbott 

2001, Aiuvalasit 2006). In general, the sediments of the Deweyville Allogroup are separated from 

younger undifferentiated Holocene fluvial and estuarine sediments by a surface created by the 

accumulation of sediment during postglacial sea level rise and local erosion surfaces created by the 

postglacial lateral migration of streams and rivers.

The Deweyville Allogroup can be further subdivided into unnamed bodies of fluvial sediments of 

alloformation rank (Blum and Vastaldo 1994). The tops of each of these allostratigraphic units are 

characterized by a single terrace, associated channel belts, sand bodies, and lateral scour surfaces. It is 

interpreted to represent a single episode of fluvial incision, stability, and lateral migration (Blum and

Vastaldo 1994, Blum et al. 2013).

The Deweyville Allogroup consists of largely coarse-grained fluvial sediments that underlie the terraces 

that lie intermediate in position between the Prairie surface and modern flood plain along the Trinity 

River. Detailed descriptions of these sediments for the Trinity River in Southeast Texas can be found in 

Garvin (2008) and Abbott (2001); for the Sabine River within the Deweyville Texas area by Bernard 

(1950); and the Brazos River by Sylvia and Galloway (2006). The Deweyville channel belt sands 

typically are coarser than Beaumont or Holocene channel belt sands. Individual subdivisions of the 

Deweyville Allogroup consist of multiple cross-cutting and therefore laterally amalgamated sandy and 
64

gravelly channel belts that contain isolated muddy paleochannel fills. Deweyville channel belt deposits 

fine upwards as is typical of point–bar deposits described from the literature, e.g., Bridge 2003. Typically, 

they are considerably coarser than corresponding point bar sediments than younger and older deposits. 

Finally, vertical accretion facies are limited to non-existent vertical accretion facies within Deweyville 

deposits (Bernard 1950, Blum et al. 1995, Abbott 2001, Blum and Aslan 2006, Garvin 2008, Blum et al. 

2013).

Little is known about systematic variations in thickness of the Deweyville Allogroup. The most detailed 

study of variations in its thickness is Garvin (2008). He found that, along the Trinity River in southeast 

Texas, the three recognized subdivisions of the Deweyville within its valley all increase in thickness 

gulfward. His “High Deweyville,” the oldest and stratigraphically lowest of these subdivisions of 

alloformation rank, increases in thickness from about 5.5 m updip to ~11 m downdip. His “Middle 

Deweyville” increases in thickness about 6 m updip to about 9–11 m downdip. His “Low Deweyville,”

the youngest and stratigraphically highest of these subdivisions, increases from about 7 m updip to about 

11 m to12 m downdip Garvin (2008).

The Deweyville allostratigraphic units have proven difficult to date. They are currently interpreted to 

have been deposited during MIS 4, 3, and 2 (Blum et al. 1995, Morton et al. 1996, Blum and Törnqvist 

2000, Blum and Aslan 2006, Blum et al. 2013). For example, of 11 samples subject to optically 

stimulated luminescence dating by Garvin (2008), eight yielded age estimates that agree with cross–

cutting relationships, whereas three yielded dates that violate cross–cutting relationships. One sample 

taken from his High Deweyville unit yielded an age representing the timing of channel belt activity from 

35 to 31 ka. Four samples from his Middle Deweyville unit yielded age ranges from 34 to 23 ka. Three 

samples from his Low Deweyville unit yielded age ranges from 23.2 to 18.8 ka. The coarse-grained 

valley fills of the Deweyville Allogroup represent the abandonment and entrenchment of valleys within 

the Beaumont alluvial plains by river systems ca. 100 ka, and multiple episodes of lateral migration, 

aggradation, and/or degradation within those valleys during the MIS 4, 3, and 2. These fluvial systems 

were graded to shorelines at midshelf or farther south (Blum et al. 1995 and 2013).

The Deweyville Allogroup is a significant stratigraphic unit in terms of both sand and cultural resources. 

The thick, laterally continuous, and sandy nature of the Deweyville channel belts make it a significant 

potential source of sand for beach restoration where they occur near the gulf bottom. As previously 

discussed, the edges of Deweyville terraces and oxbow lakes and swamps have been documented to be 

preferential locations of settlement and other activities in prehistoric times. Some of these cultural 

deposits would have been buried by eolian sediments to form stratified sites. Others would have been 

buried initially by overbank sedimentation and later by estuarine sediments as sea level rose during 

postglacial sea level rise within the valley in which they lie. Such sites depending on other factors could 

have been sheltered within a paleovalley below the level of erosion by shoreface erosion and resulting 

ravinement formation. 

The youngest of the Deweyville alloformations would have been young enough to have been 

contemporaneous, in part, with the human occupation of North America. As a result, the natural levees of 

its paleochannels and oxbow lakes are young enough to have been occupied while a particular 

paleocourse was active; this would indicate the potential for the presence of cultural resources being 

found buried within the sediments of this part of the Deweyville Allogroup.
65

6.4 The Holocene-Pleistocene Surface

Separating Prairie and Deweyville allogroups from younger terminal Pleistocene and younger Holocene 

deposits is an unconformity that has been informally referred to either as “base of the Holocene,” “top of 

the Pleistocene,” or “Holocene-Pleistocene surface.” (Figures 6.2.1.5, 6.2.1.6). The “HolocenePleistocene surface” is used in this study because in the gulfward extent, it is as old as Late Wisconsin 

(MIS 2) and is in places described as the “Late Wisconsin unconformity.” This surface is a regional 

unconformity that extends along the Gulf of Mexico coast from at least the Central Texas continental 

shelf, where it is known as the “Holocene-Pleistocene exposure surface,” across the Louisiana Coastal 

Plain and continental shelf, and eastward of the Chandeleur Islands, where it is known as the “Late 

Wisconsin unconformity.” This surface represents the former Late Pleistocene coastal plain during the 

last lowstand of sea level that occurred during the Last Glacial Maximum of MIS 2 about 29 to 11 ka 

(Heinrich et al. 2015).

Between Late MIS 2 though MIS 1, 18 to 4 ka, this low relief coastal plain was submerged as 

deglaciation led to a rapid rise in sea level. As sea level rose, the valleys of the Calcasieu, Mermentau, 

Sabine, and Trinity rivers were flooded. The lowstand flood plain soils that had developed in the terrace 

surfaces of the Deweyville Allogroup initially were inundated by rising sea level, causing the aggradation 

of a prism of fluvial and deltaic sediments that buried the terrace surfaces and formed the HolocenePleistocene surface. Later the flooding of valleys further buried this surface under estuarine deposits. 

Within paleovalleys, these fluvial deltaic deposits were in some places either partially or completely 

eroded by the transgressing shoreline within an estuary or lagoon to create a bayline or lagoonal 

ravinement surface that merged with and became the Holocene-Pleistocene surface. Within the uplands 

lying between paleovalleys and shallow paleovalleys, the shoreface of the main Gulf shoreline eroded 

deep enough to completely remove pre-existing transgressive deposits down and into either the 

underlying Beaumont Alloformation, Avoyelles alloformation, or Deweyville Allogroup. In such a case, 

the marine ravinement surface merged with and now comprises the Holocene-Pleistocene surface (Kane 

1959, Smyth et al. 1988, Thomas and Anderson 1989, Smyth 1991, Thomas 1991, Milliken et al. 2008b,

Milliken et al. 2008c, Blum et al. 2013). The nature of the Holocene-Pleistocene surface is important 

because it constrains the potential for the preservation of offshore buried and submerged cultural 

resources.

Beneath the Louisiana Chenier Plain, the Holocene-Pleistocene surface separates the Holocene 

Mermentau Alloformation from the underlying Deweyville and Prairie allogroups of the Pleistocene 

(Milliken et al. 2008b and 2008c). Gould and McFarlan (1959) inferred that all of the HolocenePleistocene surface underlying the Mermentau Alloformation from its northern pinch out at the juncture 

of the Chenier Plain with the subaerially exposed Prairie surface south to the shoreface of the modern 

shoreline is a marine ravinement surface. In contrast, Penland and Suter (1989) argue that north of Little 

Cheniere, the Holocene-Pleistocene surface beneath the Mermentau Alloformation consists of a bayline 

ravinement-flooding surface that developed behind a shoreline parallel trend of sand shoals between Ship 

Shoal and Sabine Bank. As discussed previously, it is possible that intact archaeological sites might have 

survived formation of a bayline ravinement-flooding surface and lie buried beneath the Holocene 

Mermentau Alloformation.
66

6.5 Mermentau Alloformation

The Mermentau Alloformation is an unconformity-bounded sedimentary package that fills paleovalleys 

and valleys cut by the Mermentau, Calcasieu, Sabine, Trinity, and other rivers, and their tributaries during 

MIS 3 and 4 and underlies the Louisiana Chenier Plain. The lower boundary of the Mermentau 

Alloformation is the Holocene-Pleistocene surface, which separates it from either the Deweyville 

Allogroup or Beaumont Alloformation (Figures 3.3, 6.2, 6.5, 6.7). Landward of the main gulf shoreline 

the upper boundary is either a bay bottom or the surface of the Louisiana Chenier Plain. Offshore, the 

upper boundary is an unnamed ravinement surface, which in many places separates the Mermentau from 

the overlying Sabine Bank alloformation (unnamed shelf sand in Figure 6.2). The Mermentau was first 

named as the “Mermentau Formation” by Jones et al. (1954) for water wells located near the Mermentau 

River in Calcasieu River. 
67

Figure 6.7 North-South cross-section across the Louisiana Chenier Plain showing sedimentary facies and bounding unconformities of 

the Mermentau Alloformation. 

The Holocene Pleistocene surface is the bottom bounding unconformity of the Mermentau Alloformation. Adapted from Gould and McFarlan 

(1959). Reprinted from Heinrich et al. (2015).
68

The base of the Mermentau Alloformation is the Holocene-Pleistocene surface. As previously discussed, 

it is, in part, the surface of the MIS 2 lowstand coastal plain that has been submerged by postglacial sea 

level rise. It is a composite and diachronous flooding surface that has experienced varying degrees of 

subsequent modification by ravinement. It separates the Mermentau Alloformation from underlying 

sediments and terraces of the Deweyville Allogroup and unnamed Holocene fluvial deposits, which fill 

the bottom of paleovalleys and valleys of the Calcasieu, Sabine, other coastal rivers and streams. In cores, 

the Holocene-Pleistocene surface exhibits a sharp increase in shear strength–a change from overlying soft 

gray sediments to underlying hard, orange and tan mottled mud and sand. These changes represent the 

deposition of soft gray sediments of the Mermentau Alloformation on top of sediments that have been 

altered by subaerial exposure and pedogenesis (Fisk 1948, Byrne et al. 1959, Gould and McFarlan 1959,

Milliken et al. 2008b, Milliken et al. 2008c, Heinrich 2015).

Offshore of the modern shoreline within the study area, the upper contact of the Mermentau 

Alloformation is a transgressive ravinement surface, known in this report as the “unnamed ravinement 

surface.” The unnamed ravinement surface is an erosional unconformity created by the landward 

movement of the shoreface as the result of postglacial sea level rise. This surface completely truncates the 

package of sediments of the Mermentau Alloformation underlying the Louisiana Chenier Plain along the 

modern shoreline and partially truncates paleovalley fills offshore of the modern shoreline. Outside of 

paleovalleys, the unnamed ravinement surface has completely removed the Mermentau Alloformation and 

eroded down into the Beaumont Alloformation (Siringan 1993, Siringan and Anderson 1994, Rodriguez 

et al. 2001).

As described from Texas and Louisiana estuaries, the Mermentau Alloformation includes a heterogeneous 

mixture of fine, coarse, and organic sediments. Both the onshore and offshore paleovalley fills beneath 

the Louisiana Chenier Plain consist of a variety of 1) sandy bayhead delta deposits; 2) central-basin muds; 

3) tidal inlet, tidal delta, and/or back-barrier estuarine sands; and 4) fine-grained shoreface muds. Within 

chenier plains, transgressive and regressive beach and shoreface sands are also present and form distinct 

beach ridges (Smyth et al. 1988; Thomas and Anderson 1989; Smyth 1991; Thomas 1991; Siringan 1993;

Rodriguez et al. 2001, 2004; Milliken et al. 2008b, 2008c). In the valleys of Calcasieu and other coastal 

rivers, the Mermentau Alloformation also contains a blanket of organic-rich mud that accumulated within 

cypress swamps (Nichols et al. 1996). 

Sirigan and Anderson (1994) and Rodriguez et al. 2001 and 2004) argue that the Sabine, Heald, and 

Shepard banks represent submerged paleoshorelines. They subdivided sandy sediments of these banks, 

from bottom to top, into 1) a back-barrier estuarine facies that is characterized by landward-dipping 

seismic reflectors and consists of interbedded sand and mud; 2) a fore-barrier, lower shoreface and/or 

ebb-tidal delta facies that is characterized by seaward prograding to chaotic seismic reflectors and consists 

of muddy sand; and 3) a storm reworked facies characterized by a chaotic to acoustically reverberating 

seismic reflection pattern that consists of interbedded shell hash and sand. According to their 

interpretations, the “lower back-barrier estuarine” and “fore-barrier, lower shoreface/ebb-tidal delta”

facies would belong to the Mermentau Alloformation as lying between the unnamed ravinement surface 

and Holocene-Pleistocene surface (Figures 3.3 and 6.7). They interpreted these facies to be erosional 

remnants of lower shoreface and tidal inlet sands that represent the former locations of stillstand barrier 

island trends. More recent studies, e.g., Swartz (2016) and Swartz et al. (2016), dispute this interpretation 

and argue that the unnamed ravinement surface lies directly upon Pleistocene fluvial sediments and the 

sand shoals have been reworked from them.
69

The Mermentau Alloformation is a significant unit in terms of prehistoric cultural resources. First, given 

its age, its accumulation spans the human occupation of the Northeastern Gulf Coast. Its bayhead delta 

deposits have the potential to contain in situ and intact cultural deposits. Beach ridges that are associated 

with the Louisiana Chenier Plain are associated with numerous archaeological sites. However, whether 

such sites can survive shoreface transgression remains to be seen. Second, the sediments of the 

Mermentau Alloformation in places are thick enough to protect paleosurfaces and associated prehistoric 

cultural resources from shoreface erosion where paleovalleys are deep enough. Finally, the nature of its 

lower bounding unconformity, the Holocene-Pleistocene surface, determines whether significant 

paleosurfaces are preserved well and the nature of these paleosurfaces. For example, studies of the 

stratigraphy and sedimentology of the sediments of the Louisiana Chenier Plain clearly indicate that the 

Holocene-Pleistocene surface at their base is mantled by a layer of thin, but likely intact, marsh sediment. 

This layer suggests that the paleosurface of the underlying coastal plain has suffered little, if any, erosion 

during the transgression of it by marine or bay waters. If so, then associated prehistoric cultural resources 

associated with it should be relatively intact. 

6.6 Unnamed Ravinement Surface

The unnamed ravinement surface is a transgressive ravinement surface and an erosional unconformity 

created by the landward movement of the shoreface as the result postglacial sea level rise. Offshore of the 

modern shoreline within the study area, the unnamed ravinement surface forms the upper boundary of 

either the Mermentau Alloformation, Deweyville Allogroup, or Beaumont Alloformation (Figures 6.2 and 

6.5). Along the lower shoreface, it forms the pinchout of the Mermentau Alloformation that underlies the 

Louisiana Chenier Plain and truncates paleovalleys that are filled with it and Deweyville Allogroup. For 

large parts of the study area between major paleovalleys, the unnamed ravinement surface is deep enough 

to have completely removed the Mermentau Alloformation, down into underlying Pleistocene sediments, 

and has merged with the Holocene-Pleistocene surface.

6.7 Sabine Bank Alloformation

The Sabine Bank alloformation is an informal allostratigraphic unit that, within the study area, overlies 

the unnamed ravinement surface, which forms its basal unconformity (Figures 6.2 and 6.5). The surface 

of the Sabine Bank alloformation (Ha01m in Figure 6.5) is the Gulf bottom. It is informally named from 

Sabine Bank, where it forms a large offshore shoal and has been studied in detail using seismic surveys 

and cores. Within the interfluvial areas of the inner shelf, the Sabine Bank alloformation, consists of 

offshore marine sediments that are less than 1 m thick and composed dominantly of mud. These muds 

lack storm sand beds (Siringan and Anderson 1994, Rodriguez et al. 2001 and 2004). 

As argued by Siringan and Anderson (1994) and Rodriguez et al. (2001, 2004), only the upper part of the 

Sabine, Heald, and Shepard banks lies above the unnamed ravinement surface. They interpret it as an 

upper seismic facies that is characterized by a chaotic to acoustically reverberating seismic reflection 

pattern and consists of interbedded shell hash and sand consisting of storm-reworked barrier island 

sediments. As noted above, more recent studies, e.g., Swartz (2016) and Swartz et al. (2016), dispute this 

interpretation and argue that the unnamed ravinement surface lies directly upon fluvial sediments and that 

the sand shoals consist entirely of offshore sands of uncertain provenance.
70

7. Paleogeography

7.1 Paleochannels and Their Fills

Various previous studies, including hazard surveys, in the southwest Louisiana continental shelf study 

area, have often conflated paleovalleys, paleochannels, and channel belts under a single category such as 

“paleochannels.” At various times, the same features, occasionally even in the same paper, have been 

given different designations to denote the same landform. The landforms that were found most often 

conflated are fluvial valleys, channel belts, and paleochannels. Because each of these features comes with 

its own expectations as to occurrence and preservation of cultural resources, they need to be strictly 

defined and differentiated. 

A channel can be either a subaerial or subaqueous landform. It is a geometric entity in the form of an 

open conduit through which a natural body of fluid and the sediment, which it transports, flows. On Earth, 

the fluid that occupies a channel from bank to bank is either water in a fluvial channel or an active 

turbidity flow in a submarine channel. The shape of a fluvial channel in cross section varies from a 

flattened to trapezoidal half-cylinder (Erkeling et al. 2014, NSSH 2008, Sharp and Malin 1975). A 

lengthy segment of a channel is a course, e.g., a river or stream course. A paleochannel is an abandoned, 

either historic or prehistoric, segment of river or stream that has been either partially or completely filled 

or buried by younger sediment (Neuendorf et al. 2005) (Figure 7.1 and 7.2). A continuous segment of 

significant length of an inactive historic or prehistoric segment of river or stream channel is a 

paleocourse. 
71

Figure 7.1. Nomenclature for Fluvial Valley Landforms. 

Many of the terms used as labels were copied directly from block-hazard survey maps prepared over a 

period of 40 years. Some multiple listings for features reflect the usage of different investigators working 

in different areas and/or times. Other terms, though not copied directly from hazard surveys, clearly were 

implicit in them. Terms enclosed in brackets were neither used nor implied on the hazard-survey maps 

reviewed for this investigation.
72

Figure 7.2. LiDAR DEM of Amite River Valley and Paleovalley (PV) showing channels, paleochannel, and intervening interfluves. 

Image constructed using Global Mapper and LiDAR DEMs from Atlas: The Louisiana Statewide GIS.
73

As noted by Blum et al. (2013), a number of papers, including Leopold and Maddock (1953), Wolman 

and Miller (1960), and Williams (1978), have demonstrated that the dimensions, depth and width of 

channels scale to discharge. The vertical dimension between the depth of scour and the bar top is a proxy 

for bankfull depth. In addition, the bankfull cross-sectional areas have been shown to adjust to flood 

discharges with recurrence intervals of 1 to5 years, the so-called bankfull discharge (Blum et al. 2013). 

In the Gulf of Mexico region, modern fluvial channels range from hundreds of meters in width and a few 

meters in depth to the Mississippi River. For example, the Brazos River in central Texas is approximately 

146 m wide and 17 m deep near Rosharon, Texas (USGS 08116650) and 150 m wide and 12 m deep at 

Richmond, Texas (USGS 08114000) (Heitmuller 2014). In southeast Texas, the Trinity River is about 

200 m to 400 m wide and 5 m to 14 m deep near Moss Hills, Texas; near Liberty, Texas, it is about 100 m 

to 200 m wide and 6 m to 14 m deep (Phillips et al. 2005). Along the Texas-Louisiana border, the Sabine 

River is about 125 m wide and 10 m deep near Bon Weir, Texas, (USGS 08028500). At Ruliff, Texas 

(USGS 08030500), it is about 80 m wide and 14 m deep (Heitmuller 2014). Within Louisiana, the lower 

segment of the Mermentau River had a width that varied between 75 m and 300 m and a low-water depth 

of 3 m or more. Before channelization, the lower segment of the Calcasieu River was about 40 m wide 

and 1.5 m deep (Chief of Engineers 1919). Finally, at Baton Rouge, Louisiana, the Mississippi River 

varies between 600 m and 900 m wide and between 15 m and 30 m deep. 

As reviewed extensively by Ethridge and Davidson (2011), reviewed briefly by Blum et al. (2013), and 

classified by Church (2006), modern fluvial channels exhibit a spectrum of plan-view channel patterns. 

They range between end-member patterns such as braided (rapidly shifting multi-channel systems), 

meandering (sinuous single channel), anabranching (stable multichannel systems), and straight (single 

channel) (Church 2006, Ethridge and Davidson 2011). These variations in stream pattern are interpreted 

to represent differences in stream power (proportional to the product of discharge and slope), bed material 

grain size, and the presence or absence of bank-stabilizing vegetation and/or muds (Church 2006). 

The classical model of braided “lowstand” rivers and streams was developed following the observations 

of the glacial-period lower Mississippi River by Fisk (1944) and others. However, as discussed by Blum 

et al. (2013), the concept of lowstand braided rivers and streams is only applicable in general to the 

glacial-period Mississippi River and other meltwater channels for a number of reasons. First, unlike the 

glacial-period Mississippi River, other lowstand Gulf of Mexico rivers lacked glaciation in their 

headwaters and, as a result, were not braided (Saucier 1994, Blum et al. 1995, Blum and Aslan 2006,

Blum 2007). Second, in modern, low-gradient, sedimentary basins that lack glacial influence in their 

contributing drainage areas, the braided-meandering transitions occur within uplifting to stable 

continental interiors, along rifted margins, or within a few tens of kilometers of orogenic fronts along 

active margins. As a result, in modern subsiding basins the majority of low-gradient river systems are 

dominated by either meandering or anabranching channels (Bridge 2003, Latrubesse 2008, Blum et al. 

2013). Third, both Bridge (2003) and Blum et al. (2013) argue that it is unlikely that reliable 

interpretations of paleochannel patterns can actually be made solely from 2-D seismic data, well logs, or 

outcrops. This argument is supported by interpretations of braided fluvial systems being common in 2-D

seismic, well-log, and outcrop studies; interpretations of braided channels in 3-D seismic data are rare 

(Ethridge and Davidson 2011, Blum et al. 2013). Finally, Blum et al. (2013) speculate that prehistoric

braided rivers were common in certain tectonic settings, but are uncommon in the overall stratigraphic 

record, especially after the Early to Middle Paleozoic advent of primitive land plants, which introduced 

rooting systems, and promoted the production of meandering and anabranching fluvial systems and the 

preservation of their paleochannels. Therefore, the concept of lowstand braided fluvial systems likely 

does not apply to the Texas-Louisiana continental shelf. It is likely that the majority of the paleochannels 

within this region are meandering and anabranching in nature. 
74

The assemblage of sediments that was created by the filling of a paleochannel is known as a 

“paleochannel fill” (Figures 7.1 and 7.2) It accumulates as the result of either the abandonment of a 

channel by neck cutoff, a local event that affects a single meander loop, or an avulsion, a regional event 

that affects an entire river course. Typically, channel fills are heterolithic. Normally, the upstream reach 

of a paleochannel is sand-rich because it is proximal to and remains connected with the newly active 

channel. A typical channel fill becomes increasingly mud-rich and eventually organic-rich along a 

paleochannel with distance away from where it connects to an active channel (Blum et al. 2013). Bridge 

(2003, 2006) provides a variety of detailed facies models for channel fills. Reijenstein et al. (2011) 

provide examples of paleochannels in high-resolution 2-D and 3-D seismic data. 

Because of the overall heterolithic and generally fine-grained natures of channel fills, they are typically 

unsuitable as sources of sand needed for beach restoration and similar projects. Generally, they will occur 

as sinuous ribbons of muddy sediment within a fluvial channel belt sand that is composed of sand suitable 

for restoration purposes (Figures 7.3 and 7.4). 

Figure 7.3. Hierarchy of fluvial architectural elements. 

Modified from SEPM Stratigraphy Web (2019).
75

Figure 7.4. Peveto Channel sand body on the OCS used for Holly Beach Restoration interpreted as Late Pleistocene channel belt.

Pb = point bar and Pc = paleochannel. Map from Coastal Planning and Engineering (2001).
76

7.2 Channel Belts and Deposits 

As defined by Blum et al. (2013), a channel belt is a larger scale fluvial landform that consists of the area 

of floodplain over which a fluvial channel has migrated (Figures 7.1 and 7.2). This area includes 

individual paleochannels. Similarly, according to Blum et al. (2013), a channel-belt sand body, the 

channel body of Gibling (2006), is a single story (“storey”) of sandy to gravelly sediments created by the 

lateral migration of a channel with no net aggradation (Figure 7.3). As defined and illustrated by Friend et 

al. (1979), a story is typically a flat-topped, and convex-downward body of sand, gravelly sand, to gravel 

with a major scour surface forming its base. A fluvial story typically fines upward and is overlain by 

overbank fines. The upper contact can also consist of the major scour surface at the base of an overlying 

channel belt sand body (story). Together they would comprise a multistory sand body (Potter 1963,

Friend et al. 1979, Blum et al. 2013). The thickness of a single story of a channel-belt sand body generally 

approximates the vertical distance between the maximum depth of scour and the water surface at bankfull 

stage. This thickness may be significantly greater than the observed channel depth at low flow. In seismic 

records, a single set of large-scale inclined-strata sets will represent what is observable and diagnostic at 

the scale of seismic data (Reijenstein et al. 2011, Blum et al. 2013). 

The thickness of a fluvial story and associated channel belt sand body will vary greatly with the scale of 

either the stream or river that created them. For example, in the case of the Oyster Creek course of the 

Brazos River near Dewalt, Texas, the channel belt sand body is about 15 m to 18 m thick and 2 km to 2.5 

km wide (Bernard et al. 1970). The possible range of story thicknesses is from 2 m to 3 m for small 

streams up to 30 m to 60 m, for a fluvial system the size of the Mississippi River (Blum et al. 2013). 

The channel belt and associated paleochannels are significant features in terms of both sand resources, 

paleosurfaces, and buried cultural resources. The channel belt sand body that underlies a channel belt is a 

prime target for finding economic offshore sources of sand suitable for the restoration of beaches and 

similar coastal projects. Although the bulk of the channel belt sand body may consist of suitably sandy 

sediments, any associated paleochannels will largely consist of ribbon-like bodies of muddy sediments, 

which are unsuitable for restoration projects, lying within the sandy channel belt sediments. In addition, if 

buried and preserved, the surface of a channel belt is a paleosurface and part of a relict floodplain that 

would have been in prehistoric times a favored location for prehistoric settlement. The natural levees that 

flank these paleochannels would be a preferred location for the location of prehistoric settlements and the 

accumulation of archaeological deposits (Weinstein et al. 1979a, Weinstein 1981, Gagliano et al. 1982). 

In many places, the floodplain–channel belt paleosurfaces have been destroyed by either ravinement 

processes or channel scour during the aggradation of a coastal plain or within a paleovalley during a 

marine transgression (Pearson et al. 1986).

7.3 Paleovalleys and Their Fills

The next level of fluvial landforms found within the study area are paleovalleys. Typically, the term 

“incised valley” is used in the literature instead of “paleovalley.” However, Van Wagoner et al. (1990),

Zaitlin et al. (1994), and Boyd et al. (2006) define “incised valleys” to include: 1) significant erosional 

relief with truncation of older strata, 2) juxtaposition of fluvial or estuarine strata on marine deposits, and 

3) demarcation of a significant basinward shift in facies, with subaerial exposure on interfluves. As noted 

by Blum et al. (2013), many fluvial valleys are incised, but do not meet criteria 2 and 3 of how an incised 

valley is defined. In addition, the modern analogues for “incised valleys” are typically buried beneath 

Holocene delta plains, or submerged on the shelves, and lack morphological expression. Therefore, 

following Blum and Törnqvist (2000) and Blum et al. (2013), valley or paleovalley are preferred over 

incised valley because, by definition, all valleys are incised. 
77

Like Blum and Törnqvist (2000) and Blum et al. (2013), it is assumed here that relative sea-level fall 

produces incision and an upstream-propagating wave of stream rejuvenation. Step-wise sea level fall 

results in step-wise incision and the formation of terraces and their associated allostratigraphic units. 

During sea level lowstand, a basal channel-belt sand body forms by the subaqueous scouring of migrating 

meandering or anabranching channels. Contrary to the often-reproduced illustration of the Mississippi 

River model of Fisk (1944), the bottom of a valley is never subaerially exposed. Later, a paleovalley 

eventually becomes filled by a predictable succession of sedimentary facies as sea level rises, e.g., Blum 

(1993), Boyd et al. (2006), and many other publications too numerous to mention (Figure 7.3). 

As discussed by Blum and Törnqvist (2000), Boyd et al. (2006), Blum et al. (2013), and many other 

publications, the changing fluvial and estuarine processes during a sea level cycle create a predictable 

succession of three sets of fluvial and estuarine deposits. Initially, the forced regression of the shoreline 

causes entrenchment of the river into its fluvial ridges and deltas as the river course extends seaward. If 

the river course is abandoned during this process, entrenched meanders cut into its fluvial ridge and 

beheaded distributaries of its delta are left behind as exhibited by the surfaces of the Prairie Allogroup 

units in southwest Louisiana. 

More commonly, relative sea-level not only falls to the point that deep incision of the coastal plain and 

valley formation occurs, but incision occurs in a step-wise fashion. As happened during MIS 4 through 2 

within the northwestern Gulf of Mexico coastal plain, each period of incision was separated by periods of 

lateral migration and extensive channel-belt deposition. This results in the creation of a valley where a 

series of terraces overlying channel belt sand bodies flanking either side of the valley step downward 

along strike. In this situation, the oldest channel belt sand body lies topographically higher than, and is 

cross-cut by, successively younger channel belt sand bodies. This type of degradational stacking of 

channel belt sand bodies and associated terraces is typical of the northwestern Gulf of Mexico coastal 

plain where subsidence rates are relatively low and shelf gradients are similar to the river longitudinal 

profile. As a result, major slope adjustments are not necessary as the river course extends across the shelf 

as relative sea level falls (Blum and Valastro 1994, Blum and Price 1998, Blum et al. 2013). 

Second, during a relative sea level lowstand of a glacial cycle, e.g., MIS 2, coastal rivers lay confined 

within valleys that they entrenched into the coastal plain and former highstand continental shelf. As a 

result, the shifting of either meandering or anabranching channels deposited an extensive single-story, 

channel belt sand body that overlies a scoured surface cut by lateral migration of the course within a 

coalesced channel belt. Furthermore, when relative sea level initially rose in step-wise fashion, the 

migration of a river course within the confined valley typically eroded and truncated the upper part of the 

underlying story as it deposited another channel belt sand body (Figure 7.3). This action created a 

multistoried, amalgamated, lower valley-fill consisting of superimposed channel belt sand bodies that are 

separated by laterally persistent scour surfaces. This amalgamated, lower valley-fill, following the model 

of Fisk (1944) for the Mississippi River Valley, has been commonly interpreted to mean that many 

lowstand Quaternary coastal plain rivers were “braided” during sea-level lowstands and that the 

amalgamated lower valley-fill of a paleovalley represents the braided-stream deposits (Blum 2007, Blum 

et al. 2013). The basal fluvial fill and terrace deposits represent potential sand resources if they occur 

close enough to the gulf bottom. If a paleovalleys is sufficiently deep, the terrace surfaces and basal 

channel belt within a paleovalley represent paleosurfaces that might have cultural resources preserved 

below the unnamed ravinement surface (Pearson et al. 1986). 

Finally, during rising sea level, the valley fills with a mud-rich non-amalgamated upper valley-fill 

sequence that represents aggradation in response to relative sea-level rise. In the coastal plain away from

the transgressing shoreline, the mud-rich sediments typically consist of separate fluvial channel belt, 

backswamp, and lacustrine sediments that are fundamentally avulsive and distributive (Saucier 1994,

Törnqvist et al. 2000, Blum et al. 2013) (Figure 7.3). Adjacent to the transgressing shoreline, the upper 
78

valley fill of mud-rich sediments is dominated from bottom to top, by tidal inlet and/or flood tidal delta; 

central bay mud; and upper bay and/or upper bay delta sediments. These depositional environments often 

abruptly back-step along flooding surfaces. They overlie channel belt sands at the base of river and stream 

valleys. In the case of the smaller, intrabasinal rivers, these basal fluvial sands often consist of 

multistoried, amalgamated channel belt sand bodies (Anderson et al. 2008, Milliken et al. 2008b, Milliken 

et al. 2008c). 

The transgressive part of these valley fills accumulated within a zone of backwater effects associated with 

the shoreline. As a result, it accumulated as a landward-tapering wedge that is proportional to backwater 

length. At their landward end, a trangressive valley fill onlaps onto amalgamated fluvial channel belt and 

floodplain sediments that accumulated during the MIS 4–2 period of relative sea-level fall and lowstand. 

The landward termination and onlap of trangressive valley fills migrates upstream through time as the 

valley fills (Törnqvist et al. 2001, Milliken et al. 2008b, Milliken et al. 2008c, Blum et al. 2013). 

7.4 Paleochannels Compared to Channel Celts and Paleovalleys

As previously noted, the landforms mapped in the hazard and archaeological surveys are identified by a 

bewildering multitude of names. The most common components of this nomenclature are “paleochannel”

and “channel.” However, judging from their plan view and cross-section where they cross the USGS 

seismic data, most of these paleochannels and channels are either paleovalleys or channel belts. 

Coastal plain paleovalleys differ from paleochannels in plan view, internal structure, scale, and relief. In 

plan view, paleovalleys exhibit a strong tendency to widen downstream independent of relative sea-level 

in contrast to the relatively parallel-sided and sinuous form of an observed segment of channels and 

paleochannels, e.g., Martin et al. (2011) and Reijenstein et al. (2011). Similarly, paleovalleys also differ 

noticeably in plan view from channel belts in that channel belts, of which paleochannels are a part and 

inset into, are typically of variable width. They also have roughly subparallel and often arcuate 

boundaries that differ from the edges of both paleovalleys and paleochannels. In the case of scale relative 

to that of contemporaneous channel belts, coastal-plain paleovalleys range from comparably wide to 

wider. However, both paleovalleys and channel belts are significantly wider and more variable in relative 

width than the paleochannels inset into them (Miall 2002, Ethridge and Schumm 2007, Reijenstein et al. 

2011, Blum et al. 2013). 

Paleovalleys differ from paleochannels and channel belts in cross section, as seen in seismic sections. 

Basically, the valley walls of a paleovalley extend above the height of a single-story channel belt sand. 

Thus, the fill of a paleovalley will consist of either basal single-story channel belt sand or multistory 

channel belt sand bodies overlain by a mud-rich non-amalgamated upper valley-fill of either fluvial or 

estuarine origin. The depth of incision of a channel belt will be limited to the thickness of its channel belt 

sand body. Also, the fluvial channel bar deposits of a channel belt sand body often exhibit distinctive 

inclined seismic facies, e.g., Reijenstein et al. (2011) and Blum et al. (2013). However, a paleovalley that 

has been truncated by erosion down to its basal channel belt sands would be indistinguishable from a 

channel belt sand body. Paleochannels will appear as depressions that have relief equivalent to or less

than the thickness of a single channel belt sand body. The fill of such depressions has been observed as a 

continuous and small-scale sag in seismic reflections as discussed by Reijenstein et al. (2011). 
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7.5 Southwest Outer Continental Shelf Study Area

The application of the above criteria to the southwest Outer Continental Shelf (OCS) study area indicates 

that the majority of features mapped as either channels or paleochannels by the hazard surveys are largely 

paleovalleys that in some cases are mixed in with channel belts and paleochannels. Many of the features 

exhibit the noticeably divergent, irregular plan view with distinct tributaries that are more consistent with 

paleovalleys than either channel belts or paleochannels as in West Cameron Blocks WC151, WC165, and 

WC166 (Figure 7.5 In West Cameron Blocks WC115, WC128, WC129, and WC130, the largest of these 

paleovalleys can be identified as the MIS 2 paleovalley of the Calcasieu River (Figure 7.5). A few of 

these features have the parallel sides and meandering plan view that is typical of true paleochannels as 

seen in East Cameron Blocks, EC30, EC31, and EC44. In West Cameron Blocks, WC132, WC149, 

WC150, and WC167, one large sinuous feature is large enough to be a possible channel belt (Figure 7.5). 

Unfortunately, the data are too fragmentary and inconsistent to determine the interrelationships between 

and complete extent of many of these fluvial features. Also, the lack of a representative seismic record for 

each of the specific fluvial features mapped prevented their reinterpretation. Because of this, it was 

impossible to correlate and reconstruct a relative chronology of paleovalleys, channel belts, and 

paleochannels and to confidently identify the nature of many of the mapped fluvial features. Unlike the 

Calcasieu, Sabine, and Trinity paleovalleys, it could not be determined which of these fluvial features 

either belong to the Beaumont, Deweyville, or Mermentau alloformations, because many of the fluvial 

features are truncated by the unnamed ravinement surface. As a result, they potentially vary in age from 

Late MIS 5 to either late MIS 2 or early MIS 1. 

Figure 7.5. Excerpt from plot of fluvial features mapped by offshore hazard surveys. 

Excerpt shows features interpreted to be paleovalleys (pv) and channel belts (cb). Note northwardoriented paleodrainage (npd).

However, within the hazard-survey data, there are noticeable drainage anomalies. The first is the wellknown coast-parallel paleovalleys of the Calcasieu, Mermentau, and Sabine rivers that once flowed 

westward into the Trinity River's paleovalley and then to the shelf edge. Second, possibly associated with 
80

this anomaly is a partial, northward-draining paleovalley system in West Cameron Blocks WC115, 

WC128, WC129, and WC130 (Figure 7.5). This orientation is the opposite of what one would expect,

given the gulfward slope of the present continental shelf. One possible explanation for both anomalies is 

that the crest of the forebulge of the Laurentide Ice Sheet was at one time during the Late Pleistocene 

south of the study area. As a result, a topographic east-west trough existed within the study area when it 

was an exposed coastal plain that diverted the drainage of the Mermentau and Calcasieu rivers westward 

and parallel to the coast. During the MIS 2 lowstand, this drainage pattern was superimposed on the 

exposed shelf when the Mermentau and Calcasieu rivers further entrenched into it along with northward 

draining tributaries of it. Finally, the apparent channel belt in West Cameron Blocks WC132, WC149, 

WC150, and WC167 is at right angles to and abruptly terminated by the northward-draining paleovalley 

(Figure 7.5). This suggests that the paleovalleys and other fluvial features mapped by hazard surveys 

represent an amalgamation of multiple periods of paleovalley and channel belt formation. 

With few exceptions, the relationship of the paleovalleys, channel belts, and paleochannels of the hazard 

surveys to the Pleistocene fluvial features (referred to as stream “courses,” “fluvial channels,” and “buried 

fluvial channels”) of Suter and Berryhill (1985) and Suter (1987) is problematic. From comparison of the 

fluvial features mapped by Suter and Berryhill (1985) and Suter (1987) and the hazard surveys in the 

project GIS it is clear that the Calcasieu River paleovalley closely matches one of Suter's fluvial channels 

(Figure 7.6). Nevertheless, an examination of the distribution of both sets of mapped fluvial features 

suggests that some segments of his fluvial channels likely represent highly generalized groupings of 

separate paleovalleys. However, for other of his fluvial features there is an overall lack of any definable 

correlation between the features mapped by Suter and Berryhill (1985) and Suter (1987) and those 

mapped by the hazard surveys (Figures 7.6 and 7.7). Also, from their plan and seismic characteristics 

alone, they are regarded to be neither “channels” nor “paleochannels” because they are too large, too 

deep, and often multistoried. To effectively evaluate the relationships among Suter's fluvial features and 

those mapped by the hazard surveys, Suter's analog seismic data need to scanned and reexamined in light 

of the hazard-survey mapping and new subsurface data collected. 
81

Figure 7.6. Excerpt from plot of fluvial features mapped by offshore hazard surveys and Suter 

(1987).

Excerpt shows lack of correspondence with fluvial “channels” and related features mapped by Suter 

(1987), and the disconnected nature of Hazard survey mapping.
82

Figure 7.7. Excerpt from plot of fluvial features mapped by offshore hazard surveys and Suter 

(1987).

Maps shows the lack of correspondence with fluvial “channels” and related features mapped by Suter 

(1987), and the disconnected nature of Hazard survey mapping.

Another problematic aspect of the fluvial features mapped by Suter and Berryhill (1985) and Suter (1987) 

is their relationship to onshore channel belts and paleochannels. Although clearly degraded, the 

paleochannels and channel belts exhibited by the Prairie surface can be readily reconstructed using 

historic aerial photography, USD soil surveys, and LiDAR DEMs. As a part of the preparation of 

Heinrich et al. (2002, 2003) and Heinrich (2005b, 2006), these channel belts and paleochannels were 

mapped and compiled (Figure 7.7). Comparing the onshore MIS 5 paleochannels with the fluvial features 

of Suter and Berryhill (1985) and Suter (1987) revealed that many of his fluvial features “terminate”

abruptly at the coastline without any onshore continuation, with only two observable exceptions. Given 

that his fluvial features should be falling stage, MIS 5a to MIS 3, or slightly younger, they should have 

onshore continuity with either mappable paleochannels or channel belts. One exception is the MIS 3 

Lafayette Meander Belt of the Avoyelles alloformation that should continue offshore as at least one of 

these fluvial features (channel belt or belts) offshore to the West Louisiana delta of Anderson et al. (1996) 

and Wellner et al. (2004). The other exception is the previously mentioned Calcasieu paleovalley, which 

continues onshore as Lake Calcasieu and the associated estuary. 

7.6 Paleoshorelines and Inner Shelf Shoals

Rodriguez (1999) and Rodriguez et al. (1999, 2004) collected and analyzed sediment cores, highresolution seismic data, and high-resolution side-scan sonar data from Sabine, Heald, Shepard, and 

Thomas banks on the east Texas inner continental shelf and within the current study area. They observed 

that all four shoals have similar sediment facies, structure, and presumably genesis. On the basis of the 

location of Heald, Sabine, Shepard, and Thomas banks and mapping of tidal-inlet valley fill within the 

Trinity River paleovalley, Rodriguez (1999) and Rodriguez et al. (1999, 2004) argued that the sandy 
83

sediments of the Sabine alloformation that compose these shoals are remnants of a barrier island chain 

that has been eroded and reworked by storms and wind-driven currents into offshore sand shoals. They 

argued that what they interpreted to be shoreface/ebb-tidal delta deposits are the only portion of the 

barrier island preserved and that the overlying dune, beach, and shoreface sediments were eroded and 

reworked into the upper storm-influenced facies that comprises the sand shoals. 

More recently, Swartz (2016) and Swartz et al. (2016) collected and analyzed about 90 km of high 

resolution CHIRP data that were collected by the University of Texas marine geophysics field course over 

Heald Bank and the adjacent continental shelf. They interpreted these seismic data as showing that a large 

and complex “channel” network underlies Heald Bank and the surrounding sea floor. These authors 

reported a lack of any distinct shoreface sequences associated with the bank, which they argued sits on a 

ravinement surface, the “unnamed ravinement surface” of this report, that truncates underlying layered 

channel fills of undetermined age and correlation. They propose that Heald Bank is not a drowned barrier 

island. 

8. Synthesis and Summary 

As discussed in the preceding section, the current data collections involving offshore hazard and cultural 

resources surveys have significant shortcomings in terms of data collection. One major concern is that 

they focus on what can be said to be a “postage stamp size” view that lacks integration into a regional 

framework. As a result, the absolute age, if determined at all, of fluvial and coastal landforms that are 

being mapped is determined by what is best described as the “count from the top” method. This is a 

method that has led to considerable confusion and misinterpretations in Quaternary studies elsewhere, as 

in case of the five-stage Midwest glacial chronology. Accurate dating of individual paleovalleys, channel 

belts, and paleochannels mapped by hazard and cultural resources surveys is important, because a large 

number of these fluvial landforms likely predate the human occupation and, as a result, lack any 

predictable and significant cultural resources associated with them. If their age could be established with 

any confidence, such fluvial features could be considered as potential sources of sand for coastal 

restoration.

Another concern is the confusion in the interpretation of seismic data in the identification of and 

nomenclature used to designate paleovalleys, channel belts, and paleochannels. For example, the 

examination of the available seismic data found that many paleovalleys and channel belts are being 

misidentified as paleochannels or indiscriminately designated as “paleochannels.” This is an important 

distinction, because the expected distribution of cultural resources associated within and around a fluvial 

landform will vary greatly as whether it is either a paleovalley, channel belt, or paleochannel. As a result, 

the accurate determination of the type of fluvial landform being mapped is an essential step in evaluating 

the presence of the type cultural resources present within a lease block. In case of a paleovalley that has 

been truncated below the top of its usually amalgamated, basal channel belt, it might be impossible to 

differentiate it from a channel belt. However, for cultural resource purposes, the difference between the 

two is moot because this degree of erosion of a paleovalley has destroyed any existing paleosurfaces 

within it and the potential for the existence of significant cultural resources within it.

The limited size of offshore hazard and cultural resources surveys and lack of an accessible regional 

seismic survey also presented limitations. The lack of a usable regional overview creates large data gaps 

that prevent correlation of mapping from one group of lease blocks to another and the confident and 

detailed reconstruction of drainage patterns. Also, focusing on a single lease block or group of lease 

blocks makes it easy to miss the geomorphic context of the area being evaluated. For example, it would

be hard to tell that a lease block is lying in a large paleovalley if the valley walls lie outside the immediate 

study area. A large seismic dataset was acquired for the research by Suter and Berryhill (1985) and Suter 
84

(1987). However, it exists only in hard copy format that is impractical for the vast majority of researchers 

to access and use.

Also, an overall lack of consistency was found in the interpretation of data; the manner in which they 

were mapped; and the stratigraphic and geomorphic and nomenclature used to designate features mapped. 

In the hazard and cultural resources survey report, rarely were the specific criteria used to map the 

boundaries of fluvial landforms stated. As a result, it is hard to know how the boundaries were drawn 

from the seismic data and how these differences might have influenced the mapping of these landforms. 

In rare cases, the thalweg of a fluvial landform was mapped instead of its boundaries, which made 

comparison of mapping quite difficult if not impractical. However, it does appear that in some cases effort 

was made to edge match their data with the data of older, adjacent surveys. 

Finally, in the course of this study, it was found to be very difficult to access and replicate previous 

interpretations of offshore hazard and cultural resources surveys. This could not be done because many of 

the offshore hazard and cultural resources surveys lacked a type seismic line that illustrated the seismic 

signature of each of the landforms mapped (or access to the seismic data). Similarly, the mapping 

compiled from the offshore hazard and cultural resources surveys could not be compared with the 

research by Suter and Berryhill (1985) and Suter (1987) because the bulky hard-copy format of their 

seismic lines made it impractical to reexamine critical segments of their seismic lines.

8.1 Site Preservation and Stratigraphy

Although much remains to be learned about the geoarchaeology of the Louisiana-Texas offshore 

continental shelf, some preliminary generalizations can be made about where paleosurfaces, which might 

be associated with cultural resources, can be found. First, the important surface that is associated with 

preserved cultural resources is the Holocene-Pleistocene surface. It was the Late Pleistocene, MIS 2, 

coastal plain when the northwestern Gulf of Mexico was occupied by Native Americans. The extent to 

which it is preserved and unmodified by relative sea level rise will be a major control on the preservation 

of cultural resources during the last and ongoing transgression. As previously noted by numerous other 

researchers, Gagliano et al. (1982), Pearson et al. (1986), and Waters (1992), the primary location where 

this surface will be preserved is within deeply entrenched paleovalleys. The preservation of paleosurfaces 

within paleovalleys will depend on the rate and corresponding degree of modification that they experience 

from the transgressing bayline. Also, if a paleovalley was not entrenched deeply enough into the 

surrounding coastal plain, shoreface erosion potentially could have cut deeply into the basal channel belt 

within its valley fill and largely could have destroyed any paleosurfaces that existed within it. Within the 

Pleistocene coast-parallel surfaces and associated Pleistocene and Holocene channel belts on these 

surfaces, shoreface erosion and ravinement surface formation will likely have extensively modified the 

Holocene-Pleistocene surface and any cultural deposits associated with it. 

It is theoretically possible that periods of very rapid sea level rise might have preserved strips of the 

Holocene-Pleistocene surface by the in situ drowning of existing barrier islands or chenier plains by a 

process known as step-wise retreat. In such cases, erosion of the beaches ridges during the initial 

drowning and by later tidal action would have likely heavily reworked these beach ridges and destroyed 

any associated cultural deposits. However, as the shoreface jumps to its new location, lagoonal deposits 

and a strip of the Holocene-Pleistocene surface and any associated cultural resources underlying them 

might survive the transgression.
85

In case of delta and chenier plains, there is a good chance that some cultural resources might survive the 

initial submergence along its bayline or soundline. However, the advancing shoreface and ravinement 

formation will likely extensively have eroded the surface of the delta or chenier plain and have destroyed

any associated cultural resources. More research needs to be done as detailed studies and documentation 

about how the submergence of cultural resources actually affects both surface and buried cultural 

resources is lacking.

8.2 Recommendations for Management and Protection

This research has developed a number of recommendations for management and protection.

1. In terms of the mapping of submarine paleolandforms and potential paleolandscapes: 

A. Regional terminology should be standardized.

B. There needs to be clear definition of, recognition of, and differentiation among

paleovalley, paleochannels, and channel belts.

C. A type-seismic section for specific paleovalleys, paleochannels, and channel belts 

mapped for that block should be included in the report. 

D. Geophysical data should be submitted with block survey reports for BOEM analysis

and development of regional geologic models from multiple block surveys, and 

archiving for future or alternate uses such as sand resources identification. 

E. Geologic sampling to ground-truth geophysical data and absolute dating of potential 

paleolandscapes within fluvial valleys should be conducted using appropriate 

techniques.

2. There needs to be a more regional approach to mapping submerged landforms and attendant bounding 

surfaces. This would include:

A. Block survey data (not just interpreted map products as done in this study) should be 

applied to develop a regional geologic model and conceptual model for shelf 

evolution. 

B. When new surveys are conducted, provide regional models to operators for edge 

matching with previous studies and context locally for their study area.

C. Digitize and make readily available online the USGS seismic lines of Suter and 

Berryhill (1985) and other regional surveys conducted by USGS and LGS that are 

preserved only as analog or paper forms.

D. Develop uniform classification schemes, nomenclature, and recognition criteria for 

submerged landforms and potential paleolandscapes.

3. There needs to be a better understanding of the depositional and/or erosional response during 

transgression as ravinement surfaces are produced at the shoreface and within valley estuarine 

systems to determine preservation potential of prehistoric landscapes within valley fill packages. 

4. Allow selected drilling on paleovalleys with a mandate that boring logs and “soil” samples from them 

be made available for scientific research by Quaternary geologists and geoarchaeologists. Any 

potential damage to sites will be offset by additional data that, due to funding constraints, would not 

otherwise be collected.

5. Avoidance criteria should be developed based on a strong understanding of shelf and/or valley-fill 

evolution in response to sea-level rise. The results of this study demonstrate that shelf preservation of 

paleolandscapes with potential to host prehistoric cultural resources would be extremely rare due to 
86

significant erosion that occurs during sea level rise. Regional models that identify geologic units and 

bounding surfaces based on absolute age-verification should be applied because, based on the 

synthesis presented herein, it is possible that many of the paleo-fluvial systems being avoided in the 

vicinity of the study area are far too old for potential human occupation. 
87

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