Miocene tidal-influenced sedimentation to continental Pliocene sedimentation in the forebulge–backbulge depozones of the Beni–Mamore foreland Basin (northern Bolivia)

Miocene tidal-influenced sedimentation to continental Pliocene

sedimentation in the forebulge–backbulge depozones of the Beni–Mamore

foreland Basin (northern Bolivia)

Martin Roddaz *, Stephane Brusset, Patrice Baby, Gerard Herail

IRD UR 154 LMTG-UMR 5563, Universite Paul Sabatier Toulouse 3, 14 avenue Edouard Belin 31400 Toulouse, France

Received 1 January 2004; accepted 1 November 2005

Abstract

Mio-Pliocene deposits of the forebulge–backbulge depozones of the Beni-Mamore foreland Basin indicate tidally to fluvially dominated

sedimentation. Seven facies assemblages have been recognized: FAA–FAG. FAA represents a distal bottom lake assemblage, FAB and FAD are

interpreted as tidal flat deposits, FAC and FAG are interpreted as fluvial systems, FAE sediments are deposited in a subtidal/shoreface setting, and

FAG represents a meandering fluvial system. The identification of stratigraphic surfaces (SU, MFS, and MRS) and the relationship among the

facies assemblages permit the characterization of several systems tracts: a falling-stage systems tract (FSST) followed by a lowstand systems tract

(LST), a transgressive systems tract (TST), and a highstand systems tract (HST). The FSST and LST may have been controlled by the uplift of the

Beni-Mamore forebulge, whereas TST may result from a quiescent stage in the forebulge. Subaerial unconformity two (SU2) records the passage

from a tide-influenced depositional system to a fully continental depositional system. The Miocene tidal-influenced deposits in the Beni–Mamore

Basin suggest that it experienced a connection, either with the South Atlantic Ocean or the Caribbean Sea or both.

q 2006 Elsevier Ltd. All rights reserved.

Keywords: Amazonian foreland Basin; Forebulge; Backbulge; Tidal deposits; Sequence stratigraphy; Bolivia

1. Introduction

The depositional setting of the Mio-Pliocene formations of

western Amazonia (i.e. Pebas Formation in Peru, Solimoes

Formation in Brazil) has received a lot of attention and remains

a matter of debate. Because of the scarcity and poor quality of

the outcrops, several conflicting hypotheses have been

proposed to explain the origin of the Pebas/Solimoes deposits.

One hypothesis is that they formed as a result of a catastrophic

Pleistocene flood resulting from the sudden draining of glacial

Lake Titicaca (Campbell and Frailey, 1984). Another is that the

deposits represent deltaic deposition in an enormous Pleisto-

cene–Holocene lake (Lago Amazonas, Frailey et al., 1988). A

third hypothesis suggests that they correspond to an alluvial fan

draining the Andes (Latrubesse, 1992; Latrubesse et al., 1997).

The debate has moved to arguing a fluvial versus marine origin

for the Pebas/Solimoes formations. It is now well accepted that

from the Middle Miocene to the Late Miocene, western

0895-9811/$ - see front matter q 2006 Elsevier Ltd. All rights reserved.

doi:10.1016/j.jsames.2005.11.004

* Corresponding author. Tel.: C33 5 61332599; fax: C33 5 61332560.

E-mail address: [email protected] (M. Roddaz).

Amazonia was completely flooded by a fluviolacustrine

(Hoorn, 1993, 1994a; Hoorn et al., 1995) and/or marine

(Rasanen et al., 1995, 1998; Gingras et al., 2002a,b; Hovikoski

et al., 2005) system or a combination thereof (Wesselingh

et al., 2002; Vonhof et al., 2003).

Difficulties in constraining the temporal and spatial

boundaries of the Miocene Amazonian sea/lake arose from

the lack of age and outcrop data, especially in the southern part

of the Amazon Basin. The southern extension of the Pebas

Sea/Lake has been reported as far as in southern Peru (Gingras

et al., 2002b; Hovikoski et al., 2005), whereas the northern

extension of the Paranan Sea is known to have reached the

southern part of Bolivia (Marshall et al., 1993; Hernandez et

al., 2005; Uba et al., 2005). The extent of the Pebas Sea/Lake to

the south and its possible connection with the Paranan Sea,

constituting a hypothetic Miocene Amazonian interior seaway,

is not well documented; recent studies of the Paranan and

Pebas seas (Marshall et al., 1993; Gingras et al., 2002b;

Hernandez et al., 2005; Hovikoski et al., 2005; Uba et al., 2005)

lack about 1100 km to sustain this connection (Fig. 1).

Most articles dedicated to the interpretation of the Pebas

system do not explore the role of Andean tectonics on the

evolution of the Amazonian foreland basin system. As

Journal of South American Earth Sciences 20 (2006) 351–368

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Fig. 1. Distribution of the Miocene Pebas and Paranan seas (Marshall and

Lundberg, 1996). BA, El Baul arch; CB, Chapare Buttress; FM, Freites

Formation; MAA, Monte Alegre arch; PA, Purus arch; PF, Pirabas Formation;

and VA, Vaupes arch. Southern extension of the Pebas Sea is reported in

Gingras et al. (2002b); the northern extension of the Paranan Sea is mapped in

Marshall et al. (1993).

M. Roddaz et al. / Journal of South American Earth Sciences 20 (2006) 351–368352

underlined by Paxton et al. (1996), the Miocene Amazonian

sea/lake (Pebas/Solimoes system) may be ‘a result of tectonic

mechanisms such as lithospheric loading by the Andean

orogenic wedge, or dynamic topography related to the Andean

subduction, or both.’ Furthermore, recent studies have pointed

to the influence of Andean foreland basin dynamics on the

Miocene evolution of the northwestern Amazonian paleoenvir-

onment (Hermoza et al., 2005; Roddaz et al., 2005).

The aim of this article is to present the first data relative to

tide-influenced deposits in northern Bolivia (Beni–Mamore

Basin) and document how the transition between these tide-

influenced deposits and the fully fluvial modern Amazon

drainage basin was controlled by the dynamics of the

Amazonian foreland Basin.

2. Geological setting

The western Amazon drainage basin extends from southern

Colombia to northern Bolivia. In the northern branch of the

Bolivian Orocline, the Beni–Mamore Basin forms the southern

part of the western Amazon Basin. It is limited to the southwest

by the western boundary of the Eastern Cordillera and to the

northeast by the Brazilian Craton (Figs. 2 and 3).

This foreland basin system was formed in response to

flexural subsidence driven by thrusting and loading of the

Andean retroarc thrust wedge, the forward propagation of

which has been approximately 74 km during the past 10 Ma

(Baby et al., 1997). In the structural cross-section shown in

Fig. 3, the architecture of the Beni–Mamore sedimentary Basin

is constrained by field, geophysical, and borehole data (Baby

et al., 1995, 1997). This cross-section and the geomorphic map

in Fig. 2 provide evidence of Andean tectonic control in the

Amazon Basin of Bolivia.

The Beni–Mamore sedimentary Basin can be considered the

prototype of a modern foreland system, as defined by DeCelles

and Giles (1996), in which orogenic loading leads to

partitioning into deformed and flexural provinces, such as the

wedge top, foredeep, forebulge, and backbulge depozones. The

wedge top depozone of the Beni–Mamore Basin corresponds to

the frontal part of the orogenic wedge (sub-Andean foothills,

see Fig. 2), where most of the Neogene and Pleistocene

syntectonic sediments accumulated and were progressively

deformed. The most internal syncline of the sub-Andean

foothills (Alto Beni syncline) developed as a piggyback basin

during the Neogene (Baby et al., 1995). Sub-Andean external

thrusts involved Neogene foredeep sediments. The boundary

between the present wedge top and foredeep depozones

corresponds to the sub-Andean frontal thrust. The maximum

thickness of the continental Neogene and Pleistocene deposits

filling the Beni–Mamore foredeep is approximately 6 km

(Fig. 3), located at the footwall of this frontal thrust. The

foredeep depozone consists of a 175 km wide cratonward-

tapering wedge. The forebulge is approximately 100 km wide

and poorly eroded. To the northeast, the Precambrian rocks of

the Brazilian Craton are exposed. Pleistocene sediments

accumulate between the craton and the forebulge, constituting

a 140 km wide backbulge depozone. The wedge top, foredeep,

forebulge, and backbulge zones correspond to distinct

geomorphic regions (Fig. 2). The localization of the foredeep

and forebulge depozones has been more precisely defined

through a GPS survey of longitudinal gradient (Aalto et al.,

2003).

Sedimentary deposits outcropping along the Beni River are

poorly dated. Campbell and Frailey (1984) refer to two

formations that outcrop in cutbanks along the Beni River:

Tertiary Ipururo red beds and the Pleistocene Inapari

Formation, separated by a subaerial unconformity termed the

‘Uncayali unconformity.’ Mammal-bearing faunas of ‘possible

or reputed Huayquerian age’ (Late Miocene) have also

been found within the deposits of the so-called ‘Cobija’

Formation (Marshall et al., 1993). The localities (Cobija and

Candelaria, Fig. 4) where the mammals were collected are

situated 170 and 30 km northwest, respectively, of our studied

deposits (Fig. 4).

Given that these deposits have no structural dip, reasonable

correlations can be made between our studied deposits and the

Late Miocene deposits of the Cobija Formation. Hence, the

deposits outcropping within the forebulge–backbulge depo-

zones of the Beni-Mamore foreland Basin range from Miocene

to Pleistocene.

Fig. 2. The Madeira drainage Basin (geomorphic map from USGS GTOPO-30 digital database, 1996). The dark line indicates the location of the structural cross-

section shown in Fig. 3. Modified from Baby et al. (1999).

M. Roddaz et al. / Journal of South American Earth Sciences 20 (2006) 351–368 353

3. Modern and ancient foreland deposits in the Bolivian

foreland basin system

The modern foreland basin system adjacent to the central

Andes is exemplary of a retroarc foreland basin (Jordan et al.,

1983; Jordan, 1995; Horton and DeCelles, 1997). In southern

Bolivia, this system propagated eastward since at least Late

Paleocene times (DeCelles and Horton, 2003), whereas in

northern Bolivia, the Amazonian foreland Basin started its

eastward propagation since the Neogene (Baby et al., 1997,

Fig. 3. Structural cross-section across the Beni–Mamore foreland system. Modifi

forebulge, backbulge; bounded to the west by the Eastern Cordillera and to the eas

1999). Because of its connections with the uplift of the

Altiplano, the synforeland tertiary sedimentation in the

southern Bolivian foreland basin has been extensively studied

(e.g. Horton et al., 2001; DeCelles and Horton, 2003,

references therein), whereas such studies are lacking for the

northern part. Miocene Beni deposits also should be

compared with the sedimentation and depositional environ-

ment in the ancient foreland basin systems of southern

Bolivia and the modern sedimentation of the Beni Basin

(Aalto et al., 2003).

ed from Baby et al. (1999). Note the four depozones: wedge top, foredeep,

t by the Brazilian Craton.

Fig. 4. Map of the study area. Main outcrops are black stars. Main rivers are represented. The grey filled circle represents the location of the outcrops studied by

Gingras et al. (2002b) and locations of ‘reputed’ Late Miocene mammals faunas (Marshall et al., 1993).


3.1. Ancient (Tertiary) synorogenic sediments of the southern

Bolivian foreland basin

The Tertiary synorogenic sediment succession is mainly

located in the Altiplano, the Eastern Cordillera, and the sub-

Andean zone. In the Altiplano and Eastern Cordillera, it

consists of the Santa Lucia, Potoco, Impora, Cayara, Camargo,

and Suticollo formations; in the sub-Andean zone, it is the

Petaca, Yecua, Tariquia, and Guandacay formations.

The Santa Lucia Formation (Mid-Paleocene–Middle

Eocene, Sempere et al., 1997; Horton et al., 2001) consists

of 50–300 m of quartzose, lenticular sandstone bodies

alternating with red siltstones (DeCelles and Horton, 2003)

and can be found in both the Altiplano and the Eastern

Cordillera. The sandstone bodies (1–6 m thick) are character-

ized by trough and planar cross-stratifications (lithofacies St

and Sp, Miall, 1996) and ripples. The base of the sandstone

bodies is erosive, and they typically fine upward (DeCelles

and Horton, 2003). The red siltstones are massive with

carbonate nodules, root traces, and small tubular burrows. The

Santa Lucia Formation has been interpreted as formed by

fluvial channels and overbank deposits (Horton et al., 2001;

DeCelles and Horton, 2003). The lenticular sandstone bodies

of lithofacies St and Sp are fluvial channel deposits, the

rippled sandstones are crevasse splay deposits, and the

massive red siltstones are interpreted as paleosols. The

Santa Lucia Formation ends with a massive sandy paleosol.

It is thought to have been deposited into the backbulge

depozone of the Paleogene central Andean foreland basin

system (Horton et al., 2001; DeCelles and Horton, 2003).

The Potoco Formation, only found in the Altiplano, is 3000–

6500 m thick (Horton et al., 2001). The lowermost Potoco

Formation is composed of tens of meters thick intervals of

stacked paleosols that formed during the Mid-Paleocene–

Middle Eocene. The main body (Late Eocene–Oligocene) is

composed of sheet sandstones and interbedded siltstones

(Horton et al., 2001). Sandstones are characterized by

horizontal, ripple, and trough cross-stratification; their bases

are nonerosive or slightly erosive. They are interpreted as

channel and crevasse splay deposits. The siltstones represent

overbank deposits. The Potoco Formation thus is interpreted as

deposited in a high-sediment accumulation aggradational

fluvial system. The lowermost Potoco paleosol interval records

the passage of the forebulge, and the overlying sediments

accumulated in the foredeep of the Tertiary central Andean

foreland basin system (Horton et al., 2001).

The Impora Formation (60–80 m thick) is poorly con-

strained in age and outcrop in the Eastern Cordillera. It consists

of massive gray and red mudrock with carbonate nodules,

massive nodular limestone, thin sandstone beds with roots and

trough and ripple cross-stratifications, laminated gray silt-

stones, and gray marly limestone. It is interpreted as a ‘mixture

of calcareous paleosols, lacustrine deposits and minor fluvial

deposits’ (DeCelles and Horton, 2003). The upper part of the

formation (gray marly limestone) is interpreted as profundal-

lake deposits. The formation was deposited in the forebulge to

foredeep depozones of the Tertiary central Andean foreland

basin system (DeCelles and Horton, 2003).

The Cayara Formation (190 m thick, age unconstrained) is

found in the Eastern Cordillera. It is composed of white to gray

quartzarenite sandstones with trough cross-stratifications,

termite burrows, mudcracks, and siltstone intraclastic con-

glomerates arranged in overlapping lenticular bodies (DeCelles

and Horton, 2003). Fine-grained rocks consist of pervasive

mottling siltstones with occasional nodular gypsum. The

Cayara Formation is interpreted as fluvial channel deposits,

paleosols, and playa-lake deposits (Sempere et al., 1997;

DeCelles and Horton, 2003). It is deposited in the distal

foredeep depozone of the Tertiary central Andean foreland

basin system (DeCelles and Horton, 2003).


The Camargo (O2000 m thick, Oligocene–Early Mio-

cene) and Sutillo (up to 850 m thick, Oligocene–Early

Miocene) formations are found in the Eastern Cordillera.

They are composed of (1) broad, lenticular, upward-fining,

medium-grained to conglomeratic sandstones with trough

cross-stratification (lithofacies St, Gt, Miall, 1996); (2) thick

(0.5–10 m) beds of lenticular, well-organized, imbricated

conglomerates with cross-stratifications; (3) massive mottled

red siltstones (0.25–2.5 m thick) with carbonate nodules; and

(4) thin (0.05–0.20 m) laminated siltstones (DeCelles and

Horton, 2003). The Camargo and Sutillo formations are

interpreted as large, bedload-dominated fluvial channels with

associated floodplain deposits and paleosols. They also are

interpreted as deposited in the foredeep depozone of the

Oligocene–Early Miocene central Andean foreland basin

system (DeCelles and Horton, 2003).

The Petaca Formation (Late Oligocene; Sanjinez and

Jimenez, 1975; Sempere et al., 1990) is considered to have

been deposited in arid conditions interrupted by ephemeral

stream deposits (Hermoza, 2004; Uba et al., 2005); it recently

was interpreted to have been deposited in a forebulge–

backbulge depocenter (Uba et al., 2005).

The Yecua Formation (Middle–Late Miocene, 11–8 Ma) is

found in the southern Bolivian plain. It consists of clays, black

to green marls, and subordinate calcareous sandstones

(Marshall et al., 1993). This formation bears many fossils

that indicate continental, brackish, and shallow-marine

environments (Marshall et al., 1993). The shallow marine

deposits are interpreted to record several short-lived

transgressive and regressive cycles (Uba et al., 2005). Some

authors interpret the Yecua Formation as deposited in the

Miocene backbulge depozone of the central Andean foreland

basin (Horton and DeCelles, 1997). For others, it represents

distal foredeep deposits (Uba et al., 2005).

The Tariquia Formation (Early Late Miocene) is 1400 m

thick. Its top part is considered the lateral equivalent of the

sub-Andean zone of the distal Yecua Formation (Hermoza,

2004). It is divided into two members (Hermoza, 2004). The

lower member is composed of sand bars (lithofacies Sp, St,

Sh) interbedded with reddish siltstones (lithofacies Fm, Miall,

1996). The depositional environment of this member is a

sandy fluvial meandering system (Hermoza, 2004). The upper

member consists of 20–30 m thick sand bars (lithofacies Sm,

St, Sp, Sh) interbedded with reddish siltstones (lithofacies Fm;

Miall, 1996). The upper member represents distal, sandy,

braided deposits (Hermoza, 2004). This formation is

interpreted as deposited in the Miocene foredeep of the

central Andean foreland basin system (Hermoza, 2004; Uba

et al., 2005).

The Guandacay Formation (Late Miocene–Pliocene, up to

1000 m thick, southern Bolivia sub-Andean zone) consists of

Table 1

Localization of the main outcrops. Localization of Puerto Cavinas and Riberalta ar

Name BEN 61 BEN 70 BEN 76 B

Longitude

latitude

66855.049’W 66854.516’W 66849.439’W 6

12832.243’S 12803.526’S 11842.439’S 1

a thickening, coarsening-upward sequence (Hermoza, 2004).

It is composed of medium to coarse channelized sandstones

(lithofacies St, Sh, Sp, Sm) interbedded with reddish

siltstones (lithofacies Fm) passing upward to conglomerates

(lithofacies Gmm, Gt) and thick sandstones (lithofacies Sm,

St). It is interpreted as deposited in a gravel-sand braided

system (Hermoza, 2004). The Guandacay Formation was

deposited in a piggy-back basin of the Pliocene wedge top

depozone of the central Andean foreland basin system

(Hermoza, 2004).

3.2. Modern synorogenic sediments of the northern Bolivian

foreland basin

The modern sedimentation of the Beni foreland Basin is

dominated by episodic accumulation of floodplain deposits

controlled by La Nina cycles (Aalto et al., 2003). The

episodic sediments form 20–80 cm thick packages of silts

with ‘fine horizontal lamination and other submillimeter-

scale structures and no evidence of bioturbations’ (Aalto et

al., 2003). These ‘thick’ deposits, which account for more

than 95% of the core sediment accumulation of the Beni

lowland, are interpreted as small crevasse splays formed

during levee failure (Aalto et al., 2003). Other annual regular

floodplain deposits consist of fine sediments similar to

ancient foreland deposits (see Section 3.1). These ENSO-

induced crevasse splays were the dominant mechanism of

floodplain sediment accumulation in the modern Beni

foreland Basin (Aalto et al., 2003).

4. Study area

This study is based on outcropping strata of the Neogene

deposits that occur along the Beni River (Beni province,

northern Bolivia) between Puerto Cavinas (66855.049 0W;

12832.243 0S) and Riberalta (66803.265W; 10859.091S). This

article focuses on 5 selected (among 32) exposures in

cutbanks along the Beni River (Fig. 4). The sections are

BEN 61, BEN 70 and BEN 76, BEN 82, and BEN 85 outcrops

(Fig. 4). Their localizations appear in Table 1. These sections

illustrate different parts of the sedimentary evolution of the

Puerto Cavinas–Riberalta transect (i.e. forebulge–backbulge

of the Beni-Mamore Basin).

5. Depositional environment

This study divides the strata of the Miocene outcropping

strata into seven facies assemblages. Each has been interpreted

in terms of its depositional environment.

e also given

EN 82 BEN 85 Puerto Cavinas Riberalta

6844.073’W 66821.140’W 66855.049’W 66803.265’W

1828.841’S 11821.848’S 12832.243’S 10859.091’S

Fig. 5. Representative sedimentary logs of the study area. Location of the measured sedimentary section is in Table 1.


5.1. Facies A: offshore distal bottom lake

5.1.1. Description

Facies A (40–50 cm–1 m thick, Fig. 5) comprises blue

massive silt to massive mud. Sedimentary structures are

generally absent, except for local, faint, parallel laminations.

The base of this facies, elsewhere located below the Beni

River flood level, cannot be observed.

5.1.2. Interpretation

Blue mud clays are very common in the Upper Tertiary of

western Amazonia (Pebas Formation) and indicative of a

widespread bottom dysoxia with low hydraulic energy

(Wesselingh, 2002). They have been interpreted as having

been deposited offshore in the Pebas bottom lake system, as

defined by Wesselingh et al. (2002).

5.2. Facies assemblage B: tidal flat and tidal bar

5.2.1. Description

Facies assemblage B is divided into facies B1 and B2

(Fig. 5). Facies B1 (1.5 m thick, Fig. 5) comprises regular

alternations of very fine to fine-grained sand and mud layers

exhibiting lenticular and wavy bedding (Figs. 6 and 7). The

base of this facies is channel-shaped (Fig. 8). Bidirectional,

asymmetric, bedform, heterolithic, and ripple-scale cross-

stratifications are present in wavy bedding (Figs. 6 and 7),

thus forming inclined heterolithic strata (IHS, Thomas et al.,

1987; Gingras et al., 2002b; Fig. 8). Set heights are 1.8–4.8 cm

with wavelengths of 8.5–6.5 cm. The ripples are occasionally

draped by continuous clay layers, 0.1–0.5 cm thick, containing

plant debris and coal fragments. Bidirectional currents indicate

paleoflows to the northeast and southwest. Slumps are

occasionally present, and strata coarsen upward. Facies B2

(1.5 m thick, Fig. 5) comprises massive, well-sorted, medium-

grained, yellow silty sands. Slumps are also present.


An overbank depositional environment dominated the

sedimentary filling of the ancient and modern Bolivian foreland

basin system (Horton et al., 2001; Aalto et al., 2003; DeCelles

and Horton, 2003). The overbank depositional environment

consists of levees, crevasse channels, crevasse splays, flood-

plain fines, and abandoned channels (Miall, 1996). The levee

sediments are made of rhythmically bedded units of silty,

ripple-laminated sand a few decimeters thick (lithofacies Fl;

Miall, 1996). Bioturbations or roots commonly obscure

laminations (Miall, 1996). The crevasse channel deposits

usually consist of fine- to medium-grained sandstones with

trough cross-bedding and ripple cross-lamination (lithofacies

St, Sr) forming ribbon-like bodies (Miall, 1996). Internal scours

are common (lithofacies Ss). The crevasse splay deposits are

made of fine- to medium-grained sandstones with abundant

trough cross-bedding and ripple cross-lamination (lithofacies

St, Sr) and laminated siltstones (lithofacies Fl). Plant roots and

bioturbations are common. Floodplain sediments are composed

of massive silts and muds (lithofacies Fsm), finely laminated

silts (lithofacies Fl), massive muds with desiccation cracks

(lithofacies Fm), and massive muds with roots (lithofacies Fr).

Abandoned channel fill deposits are made of the same fine-

grained facies as the preceding. The crevasse splay element is

the only one in the overbank depositional environment

composed of sandstones and mudstones, but its lobe-shaped

geometry does not account for the channel-like geometry of

facies B1 (Fig. 8). With the exception of abandoned channel fill

deposits and crevasse channels, overbank deposits have a sheet-

like sedimentary shape, inconsistent with the general channel

shape of facies B1 (Fig. 7). Abandoned channel fill deposits

commonly fine upward, which is not the case for facies B1.

Wavy and flaser bedding can also form in ephemeral streams

(Martin, 2000). However, the absence of desiccation cracks in

facies B1 excludes this mechanism for the formation of wavy

bedding. None of the facies that characterize the depositional

environments of the ancient and modern Beni foreland Basin

Fig. 6. Photograph and interpretation of wavy bedding in facies B1. Black areas represent mud and white areas sand. Dashed lines mark reactivation surfaces, and

arrows indicate the direction of the paleoflow. Pen is 15 cm long.


corresponds to facies that composed facies assemblage B.

However, IHS channel fill deposits are commonly associated

with tide-influenced point bar and tidal creeks (Smith, 1989;

Ranger and Pemberton, 1992; Rasanen et al., 1995; Falcon-

Lang, 1998; Gingras et al., 1999, 2002b; Miall, 2000). Ripple-

scale laminations with opposite bidirectional currents are

frequent in tide-dominated systems (Dalrymple et al., 1992,

2003; Lanier and Tessier, 1998). Moreover, the abundance of

wavy and lenticular bedding, reactivation surfaces, and mud

drapes on cross-bedding ripples associated with opposite

bidirectional currents in facies B1 indicate a strong tidal

influence on deposition (Nio and Yang, 1991). In modern

environments, the vertical evolution of lenticular to wavy and

flaser bedding marks an increase in the tidal range from neap to

spring (Tessier et al., 1995). Hence, this facies is interpreted as

having been deposited in a restricted tidal-flat environment.

Because facies B2 is laterally associated with facies B1, facies

B2 may represent a tidal bar. Facies assemblage B thus is

Fig. 7. Photograph and interpretation of the reversal of ripple migration in facies B1. Note plant and coal debris in the mud drapes. Black areas represent mud and

white areas sand. Dashed lines mark reactivations surfaces, and arrows indicate opposite directions of the paleoflow. Pens are 15 cm long.


interpreted as part of a tide-dominated depositional

environment.

5.3. Facies assemblage C: fluvial association

5.3.1. Description

Facies assemblage C consists of facies C1 and C2 (Fig. 5).

Facies C1 (1–2 m thick, Fig. 5) is composed of medium- to fine-

grained red sands resting over an erosonial base covered with

mud pebbles and mud breccias. Trough cross-stratification

(lithofacies St; Miall, 1996) indicates paleocurrents to the

southeast. Facies C2 (1 m thick, Fig. 5) is composed of

centimetric layers of white and red muds. Faint planar

horizontal stratifications are present. This facies also exhibits

synsedimentary normal faults (Fig. 9). Its basal contact is a

channel-shaped erosonial surface. Facies C1 and C2 are

laterally equivalent (Fig. 5).


The muds of facies C2 are probably floodplain deposits.

Because they fill a channel, they are interpreted to represent

floodplain deposits in an abandoned channel (element FACH);

Miall, 1996). The depositional environment of facies C1 can be

interpreted as channel-bottom reworking lag deposits and bank

collapse and likely corresponds to a distributary channel.

Facies assemblage C thus can be interpreted as a fluvial system

(Miall, 1996).

Fig. 8. Channel-shaped geometry of facies B1 displaying IHS (Thomas et al., 1987).


5.4. Facies assemblage D: tidal flat association

5.4.1. Description

Facies assemblage D is made up of two facies, D1 and D2

(Fig. 5). Facies D1 (1–2 m thick, Fig. 5) consists of strongly

bioturbated, fining-upward, muddy sand with plant debris

(Fig. 10). Facies D2 (1 m thick, Fig. 5) consists of alternations

of medium- to fine-grained sands and muds. The base of the

facies displays a type-B climbing ripple, containing sigmoid

laminations with an unidirectional paleoflow to the northeast.

Set heights vary from 9.8 to 13.5 cm with wavelengths of 31–

35 cm. Reactivation surfaces with mud drapes and bioturba-

tions are common (Fig. 11). The upper part of this facies is

characterized by wavy to flaser bedding laminations (Fig. 11).


Interlamination of mud, silt, and very fine-grained sand

may occur in overbank deposits (lithofacies Fl, Miall, 1996).

Their deposition results from suspension and weak traction

currents. Very small-scale ripples also may be present in

overbank sand or mud layers. Compared with modern Beni

sedimentation, laminations are thinner, and bioturbations are

present in facies D. Overall, sigmoid laminations, reactivation

surfaces, and millimetric wavy and flaser bedding alternation

of sand/mud laminae are rather rare in overbank facies. Their

occurrence, together with mud drapes in facies D2, most likely

suggests a tidal influence on the deposition. Within a tidal

environment, the dominance of one preserved current

direction implies that (1) the subordinate current velocity

was lower than the sand entrainment threshold velocity and

hence unable to transport sediment; (2) the subordinate

current velocity was capable of entraining sediment, resulting

in bedform migration, but was followed by erosion from the

subsequent stronger dominant current; or (3) no subordinate

current existed due to a dominantly fluvial influence (Davies

et al., 2003). The second mechanism constitutes probably the

best explanation, as mud drapes are preserved within the

foresets and above the reactivation surface. Wavy bedding

may be ascribed to an environment in which current velocities

and sediment supply are lower than during the formation of

flaser bedding. In modern environments, the passage from

flaser to wavy bedding ripples marks a decrease in tidal

current velocities ascribed to the passage from neap to spring

(Tessier et al., 1995). Evidence of climbing ripples combined

with the flaser-wavy bedding succession indicates deposition

from a waning flow in an intertidal setting (Dalrymple, 1992).

Facies D1 probably formed in a distributary bay environment.

Facies assemblage D thus likely represents a tidal flat

association.

5.5. Facies assemblage E: subtidal/shoreface association

5.5.1. Description

Facies assemblage E is made up of four facies: E1, E2, E3,

and E4 (Fig. 5). Facies E1 (1–1.5 m thick, Fig. 5) consists of

fine- to medium-grained sands with abundant bidirectional

trough cross-stratifications. Cross-sets, 8–30 cm in height,

show opposing current directions to the southwest and

northeast. Foresets are highlighted by clay clasts (1–3 m

thick) and occasional millimeter-thick clay drapes. Coset

packages are draped by 1–9 mm thick clays, which preserve

trough widths of 1–3 cm. This facies ends with an undulate

wavy surface with heights of 1.5–2.5 cm and wavelengths of

10–19 cm. The base of the facies is not exposed.

Fig. 9. Photograph and interpretation of facies C2 showing floodplain red fines. Note the presence of a synsedimentary normal fault.


Facies E2 (1–1.5 m thick, Fig. 5) is made up of fine- to

medium-grained yellow sands containing complex sigmoidal

cross-bedding with rare mud drapes on the foresets.

Reactivation surfaces and rare bidirectional cross-bedding

are present. These ‘megaripples’ are 33–45 cm thick.

Thicknesses of the foresets vary within a megaripple from

1 to 5 cm. Reactivation surfaces are sometimes capped

by 1–4 cm thick clay drapes. This facies ends with an

undulate wavy surface, 30 cm in wavelength and 10–15 cm

in height.

Facies E3 (1–1.5 m thick, Fig. 5) consists of fine- to

medium-grained yellow sands with unidirectional trough cross-

stratification. Trough cross-sets are 30–70 cm in height with

wavelengths varying from 1.5 to 2 m (Fig. 12). Foresets record

a current direction to the southwest. Facies E3 and E2 are

laterally equivalent.

Facies E4 (1 m thick, Fig. 5) is made up of medium-grained

sands with parallel, nearly horizontal stratification. Thicknesses

of the beds vary from 6 to 15 cm. Its contact with facies E2 is

gradual.

Fig. 10. Photograph showing horizontal laminated muds and sands in Facies D1 in the interdistributary deposits (indicated by arrow). Facies G1 and G2 constitute the

meandering modern Beni association. Person (1.7 m) for scale.



Reactivation surfaces and mud layers may occur in fluvial

point bar deposits (Collinson, 1996; Miall, 1996). During a

falling water stage, two-dimensional fluvial dunes (lithofacies

Sp, Miall, 1996) may be abandoned and exposed to subaerial

erosion, producing a curved upper dune surface. If the dune

experiences a risen water stage, it may be reactivated, and the

erosion surface may be preserved as a cross-cutting surface

within the dune (termed ‘reactivation surface’) (Collinson,

1970). Also, in low energy streams with a substantial sediment

load, mud may be deposited in the upper part of the point bar

surfaces. Because of their cohesion, these layers may not be

eroded during later high-stage flows (Collinson, 1996). Clay

clast and mud layers separating sandy foresets are rare in fluvial

point bars and likely indicate alternating, successive stages of

high and low energy. Clay clasts probably result from the

erosion of a preceding deposited mud layer. Sandy foresets can

also be bounded by reactivation surfaces. A sandy foreset

bound by either mud layers (mud couplets) or a mud layer and a

reactivation surface defines a tidal bundle (Boersma, 1969;

Visser, 1980; de Mowbray and Visser, 1984; Nio and Yang,

1991). Furthermore, thick–thin variations of the bundle, clay

drapes on the foresets, reactivation surfaces, and bidirectional

cross-bedding suggest deposition within a tidal environment.

Medium to large dune bedforms (Ashley, 1990) with

bidirectional cross-strata exhibiting facies E1 typify a subtidal

environment, with the two tidal current directions of either

similar velocity when the preserved thicknesses of the foresets

are equal or dissimilar current velocities that result in different

preserved cross-set thicknesses (Davies et al., 2003). The

undulate surface of facies E1 represents the deposition of

current ripples, defined as having a maximum height of 6 cm

and wavelength of 50 cm (Stride, 1982; Dalrymple and Rhodes,

1995) within a subtidal environment. This type of ripple can be

also superimposed on larger-scale dune bedforms (Davies et al.,

2003), like those of facies E1 and E2. Bedforms of facies E2 can

be classified as medium-sized, three-dimensional dunes

(Ashley, 1990) and are common in subtidal settings (mega-

ripples cross-bedded sets; Nio and Yang, 1991). Migration of

the megaripples (or medium-sized three-dimensional dunes;

Ashley, 1990) is mainly to the northeast, which probably

corresponds to the offshore dominant current. The wavy

undulate surface bounding the upper part of facies E2 is a

reactivation surface, indicating neap-spring tidal variations

(Nio and Yang, 1991) or seasonal fluctuations (Davies et al.,

2003). The paucity of mud drapes on foresets may indicate a

more distal setting. Facies E2 corresponds to a subtidal sand bar.

Large-scale ripples of facies E3 lie in the field of medium to

large three-dimensional dunes (Ashley, 1990) and are formed in

a subtidal setting. Parallel stratifications of facies E4 are

interpreted as the topset of the prograding sand bars of facies E2.

5.6. Facies assemblage F: meandering fluvial system

5.6.1. Description

Facies assemblage F consists of two facies: F1 and F2

(Fig. 5). Facies F1 (1–2 m thick, Fig. 5) comprises white to

yellow mud clasts in a silty matrix. The sizes of the clasts vary

from 1 to 30 cm. The base of the facies is channel shaped.

Facies F2 consists of medium- to fine-grained yellow sand with

trough cross-stratification (lithofacies St; Miall, 1996; Fig. 13)

Fig. 11. Photograph and interpretation of facies D showing ripple migration, mud drapes, and reactivation surfaces. Black areas represent mud and white areas sand.

Pen is 15 cm long.


deposited over a channel-shaped erosonial surface. It contains

mud clasts in its basal part.


Facies F1 represents basal channel-filling mass flow

deposits. The depositional environment of Facies F2 corre-

sponds to a channel. Basal mud clasts indicate reworking of

underlying facies FAC and others. Facies F1 and F2 could

represent an incipient meandering fluvial system.

5.7. Facies assemblage G: meandering modern Beni

association

5.7.1. Description

Facies assemblage G is made up of facies G1 and G2 (Fig. 5).

Facies G1 (1–1.5 m thick, Fig. 5) consists of medium- to fine-

grained sands with abundant planar cross-bedding stratifica-

tion. Facies G2 (1–2 m thick, Fig. 5) consists of red massive

muds and silts with occasional ferrolithic nodules.

Fig. 12. Photograph of facies E2 showing megaripple cross-bedded sets of opposite dips. Exposure face is normal to coset surfaces. Note wavy surface and

reactivations surface with mud drapes within facies E2. The contact between facies E2 and E4 is gradual. The subaerial unconformity (SU2), which separates facies

E4 from facies F1, is also shown. The scale is 50 cm long.



Facies G1 corresponds to lithofacies Sp (Miall, 1996), which

generally is considered to have been formed by the migration

of two-dimensional dunes. Facies G2 represents floodplain

deposits. Although this association cannot fully characterize a

fluvial assemblage, the abundance of floodplain deposits and

the rare appearance of lithofacies St suggest a depositional

system rather close to the present-day depositional system of

the Beni River. Facies assemblage G therefore is believed to

represent the modern Beni fluvial system.

6. Sequence stratigraphy

The subaerial unconformities and their correlative

conformities are used to limit sequence boundaries, and

the sequence model employed is depositional sequence IV

(Hunt and Tucker, 1992, 1995; Plint and Nummedal, 2000).

The depositional profile used (Fig. 14) is that of

Wesselingh et al., (2002). The depositional system of the

Beni–Mamore forebulge–backbulge depozones is illustrated

in Fig. 15.

Fig. 13. Photograph of trough cross-stratifications in the yellow sands of facies F2. Hammer is 30 cm long.


6.1. Subaerial unconformity (SU) and falling stage systems

tract (FSST)

Significant erosion affected the coarsening-upward tidal

flat and tidal bar association (FAB), though such erosion is

restricted to the western part of the basin and interpreted to

have formed by subaerial erosion. During the base level fall,

the SU formed coevally with the deposition of coastal

sedimentary wedges. The duration of base level fall caused

a SU to form at the top of the coastal wedges. The FAB laid

above the bottom lake sediments (FAA) with a basal scoured

surface. Above this surface, grain size increased, and

bathymetry decreased abruptly. Therefore, the coarsening-

upward coastal sedimentary wedges (FAB) are interpreted as

having been deposited as a result of a base level fall and hence

can be defined as a falling stage systems tract (FSST)

(Catuneanu, 2002). In the central part of the basin, the contact

between bottom lake sediments (FAA) and subtidal/shoreface

association (FAE) is conformable and facies change is

gradual. The FAE is interpreted as the distal equivalent of

FAB. This FSST is bound at its base by the scoured basal

Fig. 14. Depositional profile of the Pebas system (modified from Wesselingh et a

assemblage C; FAD, facies assemblage D; and FAE, facies assemblage E.

surface of forced regression (BSFR) and topped by the

subaerial unconformity (SU, Fig. 15).

6.2. Lowstand systems tract (LST) and maximum regressive

surface (MRS)

The western SU is covered by a fluvial series consisting of

fluvial channels (facies C1) and floodplain fines (facies C2).

As the FAC overlaid the SU, it recorded an incipient stage of

base level rise during normal regression when the rate of

deposition was greater than the rate of base level rise. FAC

represents a lowstand stage systems tract (LST) whose fluvial

portion has low preservation potential due to subsequent

ravinement erosion. Gliding features maintain accommodation

and preserve a low gradient for the mean depositional surface

(Catuneanu, 2002). This part of the LST is overlaid by the

tidal flat association FAD, which is sharp based and only

preserved in the BEN 70 section (Fig. 5). This sharp contact is

considered as the maximum regressive surface (MRS,

Fig. 15).

l., 2002). FAA, facies assemblage A; FAB, facies assemblage B; FAC, facies

Fig. 15. Wheeler diagram illustrating the stratigraphic architecture of the Beni–Mamore forebulge–backbulge depozones. FAA, Facies Assemblage A; FAB, Facies

Assemblage B; FAC, Facies Assemblage C; FAD, Facies Assemblage D; and FAE, Facies Assemblage E. BSFR: Basal Surface of Forced Regression; FSST: Falling

Stage Systems Tract; LST: Lowstand Systems Tract: TST: Transgressive Systems Tract; HST: Highstand Systems Tract; SU: Subaerial Unconformity; MRS:

Maximum Regressive Surface; MFS: Maximum Floading Surface. See text for explanations.


6.3. Transgressive systems tract (TST), maximum flooding

surface (MFS), and highstand systems tract (HST)

The tidal flat association (FAD) overlays the fluvial

association (FAC) and characterizes a transgression when the

rates of base level rise outpace the sedimentation rates (Plint,

1988; Posamentier et al., 1992; Catuneanu, 2002). The FAD

assemblage records a transgression and hence defines a

transgressive systems tract (TST). Its base is the previously

defined MRS of the later LST, and its top is the maximum

flooding surface (MFS), where preserved. The preservation

potential of the transgressive deposits is expected to be high

because the subsequent normal regression leads to sediment

aggradation across the entire basin (Catuneanu, 2002). The

BEN 70 section (Fig. 5) shows, in ascending order, the

succession FAC, FAD, FAC (Fig. 5b). The floodplain deposits

of FAC (facies C2) overlying the estuarine facies of FAD

(facies D2) record an increase in sedimentation rate during base

level rise, which corresponds to a late-stage base level rise

when the rate drops below the sedimentation rates, inducing

normal regression. These floodplain deposits are interpreted as

part of the poorly preserved continental. In the BEN 82 section

(Fig. 5d), the floodplain deposits of the FAC (facies C2) are

incised by the meandering fluvial system of the FAF (facies F2)

and hence belong to the HST (Fig. 15).

6.4. Subaerial unconformity 2 (SU2) and the ‘modern’

continental sequence

The surface that eroded the TST of BEN 61, the HST of

BEN 70 and BEN 82, and the LST of BEN 76 (Fig. 5)

corresponds to the erosion surface at the base of FAF and

defines a second subaerial unconformity, SU2. We do not

include in the stratigraphic architecture the meandering

modern Beni association (FAG) because of its fully continental

origin. In a fully continental setting, facies changes may be out

of phase with those driven by base level changes at the

shoreline (Miall, 1996) and thereby lead to improper

interpretations.

7. Discussion

7.1. Forebulge–backbulge dynamics

The western Amazonian Basin is considered the foreland

basin of the Cenozoic Andes. Recent studies in northern Peru

(Hermoza et al., 2005; Roddaz et al., 2005) have shown that the

north Amazonian foreland basin entered a four-component

configuration (wedge top, foredeep, forebulge, and backbulge

depozones) after the Middle–Late Miocene (15–10 Ma). In

southern Bolivia, the foreland basin system initiated at least in

Paleocene times (DeCelles and Horton, 2003). The area studied

herein comprises the present-day forebulge and backbulge

depozones (Baby et al., 1999; Aalto et al., 2003). Depending on

the rheology of the lithosphere and the rate of thrust-wedge

advance, the forebulge can migrate toward the orogen (Patton

and O’Connor, 1988) or the craton (Price and Hatcher, 1983;

Currie, 1997; Horton et al., 2001; White et al., 2002; DeCelles

and Horton, 2003) or remain stationary (Waschbusch and

Royden, 1992). The wedge front of the Beni–Mamore foreland

basin has migrated cratonward, recording a 74 km shortening

since the Late Miocene (10 Ma; Baby et al., 1997).

Depositional architecture in the forebulge–backbulge

depozones depends on base level rise and fall, which is

controlled by the positive or negative subsidence of the

forebulge with respect to the stationary craton, as well as

eustatic fluctuations. Falling stage systems tracts result from

forced regression imposed by base level falls. In a foreland

basin setting, this finding offers evidence that stratigraphic

sequences are correlated with local to regional tectonics (e.g.

Miall, 2000). In the study area, an eustasy-induced base level

fall would have generated an FSST in the western forebulge

and eastern cratonic parts of the basin. The FSST is only

located in the forebulge area. Therefore, it is likely that the

FSST was created by local/regional tectonics. Two mechan-

isms capable of generating uplift in this part of the foreland

basin may be envisaged: fault-related or forebulge flexural

uplift. Faulting would have induced steep topographic

gradients unable to preserve the FAB association of the

FSST. Slumps and soft sediment deformation are common in


tidal estuary or delta inlets but not in the FAB. Their

occurrence in the FAB indicates slope creation during

sedimentation. However, because the FAD association is

preserved, the slope cannot have been greater than approxi-

mately 18, and the sloping rate must have been low. This trend

is compatible with incipient forebulge uplift, which is likely to

have controlled the deposition of the FSST. The succeeding

systems tracts (LST1, TST, and HST) were deposited during

base level rise but depend only on sedimentation. During

LST1, the sedimentation rate is assumed to have been high, and

sediment was supplied by the erosion of the adjacent Andes.

Once a steady state was reached, the sediment was trapped in

the eastern foredeep. In the forebulge–backbulge, the sedi-

mentation rate drops, and TST can develop. The subsequent

HST record an increase in sediment supply, which can be

correlated with renewed erosion within the Andes to generate

normal regression. Such an increase in supply can be ascribed

to thrust activity renewal, whose climax was later recorded by

eastward forebulge migration, the occurrence of SU2, and

emplacement of the modern Beni fluvial system.

7.2. Paleoenvironment implications

The Miocene depositional environment of western Amazo-

nia (Pebas/Solimoes) is still a matter of debate. It is unclear

whether a fluviolacustrine system (Hoorn, 1993, 1994a,b;

Hoorn et al., 1995), a tidally influenced sea (Rasanen et al.,

1995; Gingras et al., 2002a,b; Hovikoski et al., 2005) or tide-

influenced megalake (Wesselingh et al., 2002; Vonhof et al.,

2003) occupied Amazonia. Controversy comes from the

freshwater signature of mollusks found in the tidal deposits

of the Pebas system (Vonhof, 2003). Tidal deposits can also

form in freshwater environments (Kvale and Archer, 1990;

Dalrymple, 1992; Shanley et al., 1992; Greb, 1993; Archer

et al., 1994; Eberth, 1996; Kvale and Mastalerz, 1998). This

could be due to (1) inland propagation of tide influence (e.g.

Saint Lawrence river in Canada; Kvale and Mastalerz, 1998),

(2) extension of riverine freshwater plumes offshore in a tide-

dominated area (Amazon River; Lentz and Limeburner, 1995),

or (3) high runoff caused by high rainfall in a tropical climate

(Kvale and Mastalerz, 1998). A combination of these could be

a possible explanation for the freshwater signature of the Pebas

mollusk fauna.

The tide-influenced deposits of the Beni Basin necessitate

an open connection with the Atlantic Ocean during the Late

Miocene. In the Late Miocene, the Paranan transgression

reached at least southern Bolivia (Marshall et al., 1993). Recent

studies in the Brazilian Acre Basin (Rasanen et al., 1995;

Gingras et al., 2002b) and southern Peru (Roddaz et al., 2004;

Hovikoski et al., 2005) report Miocene tidal deposits with

mesohaline ichnofaunas (Gingras et al., 2002b; Hovikoski et

al., 2005). Given the poor stratigraphic control on the age of the

Beni tidal-influenced deposits, they could represent either the

southern extension of the tidal deposits of the Acre Basin or the

northern extension of the marine Yecua Formation. Another

possibility is that they record the presence of the putative

Miocene Amazon Seaway (Rasanen et al., 1995; Webb, 1995).

8. Conclusions

This study shows that the Miocene deposits of the

forebulge–backbulge foreland Beni–Mamore Basin were

tidally influenced and that the sedimentary evolution from

tidal to fully continental deposits was controlled by the

forebulge–backbulge dynamics of the Beni–Mamore foreland

basin system. Given the poor stratigraphic control on these

deposits, they could be the northernmost deposits of the

Miocene marine incursions within the Pebas system, the

southernmost deposits of the Late Miocene Paranan ingression,

or the central deposits of the Miocene Amazon interior seaway.

More work should be carried out to establish a more precise

spatiotemporal framework for these tidal deposits.

Acknowledgements

This research was supported by IRD, INSU grant

99PNSE59 (Tectonique, erosion et sedimentation dans le

bassin de l’Amazone: du Mio-Pliocene a l’Actuel) and INSU

grant «Erosion des Andes». We thank Jose Darrozes and

Magali Strub for field assistance. We also thank Pierre-Olivier

Antoine, Frederic Christophoul, Jose Darrozes, and Jean-

Claude Soula for reviewing previous versions of this article.

Mike Leeder and an anonymous reviewer offered helpful

reviews that contributed to greatly improve this article.

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