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Clement, Alastair, Fuller, Ian, & Sloss, Craig(2017)Facies architecture, morphostratigraphy, and sedimentary evolution of arapidly-infilled Holocene incised-valley estuary: The lower Manawatu val-ley, North Island New Zealand.Marine Geology, 390, pp. 214-233.

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https://doi.org/10.1016/j.margeo.2017.06.011

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Facies architecture, morphostratigraphy, and sedimentaryevolution of a rapidly-infilled Holocene incised-valley estuary:The lower Manawatu valley, North Island New Zealand

Alastair J.H. Clement, Ian C. Fuller, Craig R. Sloss

PII: S0025-3227(17)30192-5DOI: doi: 10.1016/j.margeo.2017.06.011Reference: MARGO 5639

To appear in: Marine Geology

Received date: 26 April 2017Revised date: 20 June 2017Accepted date: 28 June 2017

Please cite this article as: Alastair J.H. Clement, Ian C. Fuller, Craig R. Sloss , Faciesarchitecture, morphostratigraphy, and sedimentary evolution of a rapidly-infilled Holoceneincised-valley estuary: The lower Manawatu valley, North Island New Zealand, MarineGeology (2017), doi: 10.1016/j.margeo.2017.06.011

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Facies architecture, morphostratigraphy, and sedimentary evolution of a

rapidly-infilled Holocene incised-valley estuary: the lower Manawatu valley,

North Island New Zealand

Alastair J.H. Clement1*, Ian C. Fuller1,2, Craig. R. Sloss3

1 Physical Geography Group, Institute of Agriculture and Environment, Massey

University, Private Bag 11-222, Palmerston North 4442, New Zealand

2 Innovative River Solutions, Institute of Agriculture and Environment, Massey

University, Private Bag 11-222, Palmerston North 4442, New Zealand

3 School of Earth, Environmental and Biological Sciences, Queensland University

of Technology, GPO Box 2434, Brisbane, QLD, 4001, Australia

* Corresponding author: Tel: +64 6 951 7847; Email: a.clement@massey.ac.nz

Abstract

Water well logs and vibracores recovered from the lower Manawatu valley record

the facies architecture, morphostratigraphy, and sedimentary evolution of this

drowned river incised-valley estuary over the past c. 30,000 years. The Manawatu

incised-valley estuary was rapidly infilled during the Holocene, and as a result

preserves a complete record of Holocene estuary evolution. The base of the

Holocene fill is marked by fluvial gravels deposited by the Manawatu River during

the lowstand of the Last Glacial Maximum. The gravels are overlain by back-

stepping transgressive fluvial silts and clays. The transgressive fluvial unit is

present only in the lower reaches of the valley due to accelerated sea-level rise

during the latter stages of the marine transgression. Large volumes of

transgressive marine sediment were delivered to the mouth of the valley by rising

sea-levels during the early Holocene, forming a large subaqueous tidal delta. This

delta accreted and subsequently evolved to form a subaerial barrier in the mid-

Holocene, as sea-levels stabilised during the sea-level highstand. The central basin

of the estuary was rapidly infilled by marine and fluvial sediment, and

consequently lacks the mud facies characteristic of less-rapidly-evolved estuaries.

The transition from the sandy central basin estuarine environment to the

overlying floodplain environment was extremely rapid. In some locations in the

valley the transitional fluvial bay-head delta is missing completely, while in most

locations it is present only as a very thin deposit, indicative of this rapid

environmental transition. Infilling of the estuarine central basin was extremely

rapid, taking only 2,000-2,700 years. The unique features of the Manawatu

estuary’s evolutionary history are interpreted to primarily reflect the rapid infilling

of the palaeo-estuary driven by the high rates of sediment supplied to the valley

from both marine and fluvial sources.

Highlights

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High fluvial and marine sediment delivery drove rapid infilling of the

estuary

The Manawatu estuary was completely infilled within 2700 years

The estuary lacks central basin muds and a bay-head delta facies

The Manawatu contrasts with estuaries with more limited rates of

sediment input

Keywords

Holocene, incised valley, estuary, sea-level rise, accommodation space,

morphostratigraphy, New Zealand

Introduction

The late Quaternary sedimentary fill of incised-valley estuaries preserves a nearly

complete record of the sedimentary and geomorphic evolution of these coastal

environments in response to changes in sea-level, accommodation space,

sediment supply, and tectonics (e.g., Dalrymple et al., 1992, 1994; Boyd et al.,

2006). The sedimentary record of incised-valley estuaries is complex, reflecting

the initial formation of accommodation space by fluvial incision during periods of

lower sea-level, followed by subsequent infilling through the interaction of fluvial,

wave, and tidal processes during sea-level rise and the following sea-level

highstand (e.g., Zaitlin et al., 1994; Boyd et al., 2006). The strong relationship

between the relative influence of fluvial, wave, and tidal processes, and the

morphostratigraphy of incised-valley estuaries (e.g., Bird, 1967; Klum and Byrne,

1967; Roy et al., 1980, 1984a; Nicholl, 1991; Reinson, 1992; Dalrymple, 1992) was

distilled into two distinctive end-member inter-gradational tripartite facies

models of wave-dominated estuaries and tide-dominated estuaries (e.g.,

Dalrymple and Zaitlin, 1989; Dalrymple et al., 1992). Integration of these facies

models with principles of sequence stratigraphy (e.g., Allen and Posamentier,

1993, 1994; Ashley and Sheridan, 1994; Nichol et al., 1994; Zailtin et al, 1994)

formalised the geological definition of an estuary as a transgressive coastal

environment that receives sediments from both fluvial and marine sources, and

occupies the seaward reaches of an incised valley (e.g., Dalrymple et al., 1992;

Zailtin et al., 1994; Boyd et al., 2006).

Subsequent application of these models of incised-valley estuaries focused on

elucidating how variations in boundary conditions such as hydrodynamics,

sediment supply, accommodation space, tectonics, and antecedent

geomorphology influenced the morphostratigraphy and evolution of estuaries

that deviated from the classic types (e.g., Heap and Nichol, 1997; Dillenburg et al.,

2000; Cooper, 2001; Lobo et al., 2003; Fitzgerald et al., 2005; Sloss et al., 2005,

2006a, 2006b, 2007, 2010; Switzer et al., 2009; Burningham, 2008; Wilson et al.,

2007; Abrahim et al., 2008; Brothers et al., 2008; Kennedy et al., 2008; Chaumillon

et al., 2010; Clement et al., 2010; Peterson and Phipps, 2016; Gregoire et al.,

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2017). This attention was apt, as the case studies from which the models of

incised-valley estuaries were distilled were drawn from large estuaries (40-100 km

length), with sizable catchments (~13,000-85,000 km2), on tectonically stable

coastlines, with comparatively low river discharges and limited sediment supplies

(Wilson et al., 2007; Clement et al., 2010).

The majority of studies of Holocene incised-valley estuaries have focused on

systems in only the early or intermediate stages of sedimentary infilling. In

contrast, the Manawatu estuary, with a high sediment supply, relatively small

catchment area, comparatively large discharge, and situated in a tectonically-

active landscape, represents a completely infilled incised-valley estuary, now

dominated by sediment bypass and coastal progradation. Accordingly, this study

provides a complete record of early-stage infill through to a mature incised-valley

estuary, extending previously constructed geomorphological and sedimentary

evolutionary conceptual models (e.g. Dalrymple et al., 1992; Roy et al., 1980; Roy,

1984; Sloss et al., 2006a, b, 2007), and provides new insights into the influence of

systems controls on the facies architecture, morphostratigraphy, and sedimentary

evolution of rapidly-infilled wave-dominated incised-valley estuaries.

Site description

The lower Manawatu valley is a broad coastal plain representing an infilled

incised-valley system located on the southwest coast of the North Island, New

Zealand (Fig. 1). The North Island is positioned on the Australian tectonic plate,

which is being underthrust from the east by the subducting Pacific Plate at a rate

of 40-50 mm a-1 (Walcott, 1978; Lamb and Vella, 1987; Cashman et al., 1992).

Much of the strain of the plate collision is being transferred to the leading edge of

the Australia plate. As a result the North Island landmass is undergoing crustal

shortening and is actively deforming. The eastern margin of the coastal plain is

bounded by the Tararua-Ruahine Ranges, part of the axial range of mountains

that extend throughout the North Island (Fig. 1). The ranges are essentially a

series of horst blocks (Marden, 1984) that have been uplifted over the past 1-2

Ma (Lamb and Vella, 1987; Heerdegen & Shepherd, 1992). Uplift rates vary along

the length of the ranges from 1-4 mm yr-1 (e.g., Wellman, 1972; Ghani, 1978;

Pillans, 1986; Whitehouse and Pearce, 1992), as the constituent fault blocks have

varying deformational histories (Heerdegen and Shepherd, 1992; Kamp, 1992a,b).

West of the axial ranges lies the Wanganui Basin, a 200 by 200 km ovoid back-arc

subsiding sedimentary basin. Continental shortening across the basin has resulted

in the formation of a number of northeast striking active reverse faults, many of

which are buried and fail to penetrate the surface (Begg et al., 2005). Where the

faults are buried at depth they manifest anticlinal folds in the landscape (e.g.,

Jackson et al., 1998; Clement and Brook, 2008). The southern margin of the

coastal plain is marked by a northeast trending 2-2.5 km wide elongate basement

greywacke fault-bounded horst block known as the Poroutawhao High (Fig. 1).

The upper surface of the block has been uplifted to c. 1.5-8 m above present

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mean sea-level (PMSL; Te Punga, 1953; Hesp, 1975; Brown, 1978, 1984; Stevens,

1990).

Remnants of a Last Interglacial (Marine Isotope Stage 5e, 128,000-116,000 years

BP) marine terrace are preserved along the western flanks of the axial ranges

adjacent to the coastal plain (Fig. 1). The terrace surface rises from 30 m above

PMSL in the south, to over 90 m above PMSL in the north, reflecting a northerly

increase in uplift rates along the ranges. The terrace is also tilted laterally towards

the coastal plain, implying that uplift rates are lower on the plain than in the

ranges (Heerdegen and Shepherd, 1992; Uemura and Shepherd, 2006). The

marine terrace has been dissected by a number of streams flowing out of the

ranges and onto the coastal plain, forming a series of box-shaped valleys filled

with Holocene marine and alluvial sediments (Cotton, 1918; Hesp and Shepherd,

1978). At the northern end of the marine terrace the Manawatu River emerges

from the ranges through the Manawatu Gorge, an antecedent cutting maintained

by the Manawatu River. The course of the river across the coastal plain has been

influenced by the presence of the Pohangina and Himitangi anticlines and the

Poroutawhao High. The Pohangina Anticline has directed the Manawatu River to

the southwest when it emerges from the gorge, while the Himitangi Anticline and

Poroutawhao High have anchored the mouth of the river at Foxton (Hesp, 1975;

Stevens, 1990; Heerdegen and Shepherd, 1992).

A flight of three well-defined late-Pleistocene fluvial aggradation terraces

corresponding to MIS 4, 3, and 2 are preserved in the upper part of the valley

between the Manawatu Gorge and Palmerston North City (Fig. 1). This reach of

the river channel has a relatively steep gradient (0.0012), is gravel-bedded, and

has a sinuosity of 1.4 (Page and Heerdegen, 1985; Clement et al., 2010). A series

of palaeo-meanders are preserved in the Holocene and modern alluvium of the

upper valley, indicating that the sinuosity of the reach has recently changed (Page

and Heerdegen, 1985; Clement et al., 2010; LoRe et al., in review). In the lower

valley, below Palmerston North City, the late-Pleistocene aggradation terraces are

buried beneath the Holocene coastal plain (Hesp and Shepherd, 1978; Heerdegen

and Shepherd, 1992; Begg et al., 2005). From Opiki to the ocean the channel has a

sinuosity of 2.4, while the floodplain has a low gradient (0.0002). As a

consequence the river is no longer competent to transport gravel, and becomes

sand-bedded. At Opiki the Manawatu River is joined on the true right by the

Oroua River (Fig. 1). The Oroua floodplain also features MIS 2 fluvial aggradation

terraces which plunge beneath the Holocene sedimentary fill (Begg et al., 2005).

The lower Manawatu River drains a catchment of ~5,950 km2, has a mean annual

discharge of 102 m3 sec-1, and a mean annual flood of 1450 m3 sec-1 (Duncan,

1992). Floods of up to 4,500 m3 sec-1 have been recorded (Clement et al., 2010).

The current annual sediment yield carried by the lower Manawatu is estimated at

~3.7 Mt yr-1 (Hicks et al., 2011).

Coastal progradation and dune building were initiated on the Manawatu coast

following the culmination of the Holocene marine transgression. Both were

initially fed by sand that had migrated landward across the continental shelf with

rising sea levels, and were sustained by the sizable marine sediment supply

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delivered to the Manawatu coast by southward longshore drift from rivers

discharging to the north (e.g., Gibb, 1978; Clement et al., 2010). The Manawatu

coast is particularly suited for transgressive parabolic dune formation, as the coast

has a favourable wind and wave regime, a large supply of fine-grained sand, and

the beaches typically feature wide, gently sloping foreshores forming large source

areas for sediment mobilisation and aeolian transport (Clement et al., 2010).

Three phases of Holocene dune building have taken place (Fig. 1; Shepherd &

Price, 1990; Muckersie and Shepherd, 1995): the Foxton phase, from c. 7,700 cal.

yr BP to c. 1,600 cal. yr BP; the Motuiti phase, from c. 1,000 cal. yr BP to c. 500 cal.

yr BP; and the Waitarere phase, less than 500 years old. The result is the largest

transgressive dune field in New Zealand, covering c. 900 km2.

The Manawatu coast is wave-dominated (Clement et al., 2010) and on the cusp of

being mesotidal, with a spring tidal range of ~2.2 metres and a neap tidal range of

~0.9 m (Clement, 2011). Waves generated in the northern Cook Strait area

approach the Manawatu coast from the southwest, with a maximum fetch of

~100 km, resulting in moderate wave energy. Waves approaching the coast from

the Tasman Sea to the west have a much greater fetch, resulting in a dominant

westerly swell (Clement et al., 2010). It has been widely suggested that during the

Holocene marine transgression the lower Manawatu valley was inundated by

rising sea-levels, forming an expansive estuary protected from open marine

influences by the Porotowhao High and Himitangi Anticline (Fig. 1; e.g., Te Punga,

1953; Rich, 1959; Hesp, 1975; Hesp and Shepherd, 1978; Heerdegen and

Shepherd, 1992; Clement et al., 2010). The estuary is inferred to have extended

inland as far as Opiki (e.g., Rich, 1959; Fair, 1968; Hesp and Shepherd, 1978;

Heerdegen and Shepherd, 1992), or Palmerston North City (Fig. 1; Hesp, 1975),

approximately 30 – 40 km from the present-day coastline.

Today the estuary covers an area of approximately 2 km2, and lies approximately

4 km seaward of the position of coastline at the culmination of the most recent

post-glacial marine transgression c. 7,500 cal. yr BP (e.g., Rich 1959; Shepherd et

al., 1986; Heerdegen and Shepherd, 1992; Clement et al., 2016) due to coastal

progradation (Fig. 1). Total progradation of the coast is estimated at ~4 km, at an

average rate of 0.6-0.8 m per year (e.g., Shepherd et al., 1986; Heerdegen and

Shepherd, 1992). This suggests an extremely rapid and significant rate of infilling

compared to international examples of wave-dominated incised-valley estuaries.

The Manawatu valley therefore presents a unique opportunity to compare and

contrast the facies architecture, morphostratigraphy, and sedimentary evolution

of a rapidly-infilled system with the relatively immature models constructed from

comparatively less ‘active’ coastal landscapes.

Methods

Vibracore collection and analysis

The sedimentary infill and geological evolution of the lower Manawatu valley was

reconstructed using the facies associations of 65 largely undisturbed cores

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recovered using a vibracorer driven by an electric concrete vibrator. The

vibracores were recovered from open drains first excavated in the late nineteenth

century by settlers to the valley to convert swamps into arable land. Vibracoring in

the drains was advantageous as it ensured that the substrate being cored was

fully saturated, and meant that with the logging and sampling of the drain sides

longer sedimentary sections could be analysed than would have been possible if

cores had been taken from the surface of the floodplain. The retrieved vibracores

were opened and visually logged for colour, sediment texture, lithological

composition, the presence of macrofossils, and significant facies changes. In

selected vibracores approximately 50 g (wet weight) of sediment was collected

every 10 cm, as well as at locations of significant facies changes, for sediment

analysis. Sediment analysis was conducted on selected cores to determine

sediment constituents, textural characteristics, and to distinguish between

sediments from different depositional environments. The vibracores were

adjusted for compaction or expansion (measured versus recovered length). The

position and elevation of the vibracores was determined through RTK-dGPS

surveys.

Water well logs analysis

Stratigraphic data obtained from 387 water wells drilled across the Manawatu

coastal plain were also used to reconstruct broad-scale features of the

sedimentary infill of the lower Manawatu valley. The data is maintained by the

regional authority (Horizons Regional Council) as part of permit requirements for

the drilling of water wells. The well logs complement the vibracores, as while the

vibracores could be studied in fine detail, they did not reach depths greater than -

4 m below PMSL. In contrast, the well logs reach depths of up to 290 m below

PMSL. However, the degree of detail in the water well logs is variable as there is

no standard or uniform scheme for lithological descriptions to which water well

drillers must adhere. Accordingly, generic lithological descriptions such as 'sand'

or 'gravel' are common. Colours, such as ‘blue sand’ or ‘brown gravel’ are

frequently mentioned. Seemingly separate lithological units are frequently

compressed into a single description, such as ‘sand and clay’. The intent of such

descriptions (sandy clay, clayey sand, or interbedded sands and clays) is often

unclear. Brown's (1990) guide for well drillers logging water wells was used to

interpret driller's descriptions. The stratigraphic data recorded in the well logs is

considered sufficient to reconstruct only broad-scale features of the sedimentary

infill of the valley. Distinguishing between genetically similar lithological units, for

example transgressive sand sheets, barrier sands, dune sands, sand flats and

fluvial sands, is not necessarily possible simply by a description of ‘sand’.

However, it is possible to assess the broader significance of genetically different

sedimentary units: gravel vs sand, vs silt, vs clay. Further indicators of depositional

environment include the presence or absence of shells, wood, peat, and other

organic material. It is also possible to distinguish between greywacke bedrock and

sandstone or mudstone (recorded as ‘papa’) based on driller’s descriptions.

Interpretations were made using the original driller's descriptions; no

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standardisation of the driller's descriptions was attempted (cf., Schumacher,

1999). Pictorial logs were created following the interpretations.

Cross-section construction

Cross sections through the valley fill were constructed using ArcGIS Desktop

10.3.1 using the eXacto cross-section tool (Carrell, 2009). Final detailing and

presentation of the cross-sections was performed in Inkscape 0.92. Where there

was coverage, the surface profile for the cross-sections was generated using a

LiDAR-derived 1 m DEM supplied by Horizons Regional Council. For small areas

outside the LiDAR coverage the surface profile was generated using a subset of

the ECOSAT 15 m DEM of New Zealand. Surface elevations were sampled at 20 m

intervals along each profile line. Elevations for the water well cross-sections were

exaggerated by a factor of 70-times; elevations for the vibracores cross-sections

were exaggerated by a factor of 500-times.

Radiocarbon age determinations

A single sample of wood and 13 marine mollusc samples (Austrovenus stutchburyi)

recovered from the vibracores were radiocarbon dated using accelerator mass

spectrometry (AMS) at the Waikato Radiocarbon Dating Laboratory, Waikato

University (Tab. 1). A number of published and unpublished radiocarbon age

determinations from wood and shell samples were also utilised in the facies

reconstruction (Tab. 1). Conventional radiocarbon ages were calibrated to sidereal

(calendar) years (expressed as cal. yr BP) using the radiocarbon calibration

program CALIB REV 5.0.1 (Stuiver and Braziunas, 1993; Stuiver and Reimer, 1993).

The ages of wood samples were calibrated using the non-marine calibration curve

for the Southern Hemisphere (SHCal13; Hogg et al., 2013). The age

determinations of the A. stutchburyi samples were calibrated using the marine

model calibration curve Marine09 (Reimer et al., 2009) with a ΔR value of -7 ± 45

to correct for the marine reservoir effect (e.g., Smith and James-Lee, 2009;

Hayward et al., 2010a, b; Goff et al., 2010; Clement, 2011; Tribe and Kennedy,

2010; Clement et al., 2016). All calibrated radiocarbon ages are presented using

the 2-sigma uncertainty term (95.4 per cent degree of confidence).

Results

Water-well log facies

Cross-sections constructed using the water well logs were produced for the valley

upstream of Opiki (Figs 2 and 3), and the lower valley downstream of Opiki (Figs 2

and 4). Eight individual facies were identified in the water-well logs based on: (i)

drillers’ descriptions of lithology, and their identification of shells, wood, and

other organic content; (ii) the relationships between lithological units in adjacent

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water wells, and; (iii) the valley geomorphology and surface geological mapping.

The characteristics of each facies are summarised in Table 2.

Fluvial gravels (FG)

A laterally-extensive gravel unit dips in a southwesterly-to-westerly direction

down the length of the valley. This unit is recorded in the well logs with

descriptions such as ‘gravel’, ‘metal’, or ‘sandy metal’. The gravel unit is

interpreted as the Last Glacial Maximum (LGM; MIS 2) fluvial aggradation surface

of the Manawatu River, referred to locally as the Ashhurst or Ohakean gravel (e.g.,

Fair, 1968; Heerdegen and Shepherd 1992; Milne, 1973; Clement and Fuller, 2007;

Litchfield, 2003). LGM fluvial aggradation terraces comprising this gravel are

preserved in the valley upstream of Palmerton North City (Fig. 1), and it has been

hypothesised that these terraces plunge beneath the Holocene fill of the lower

valley (e.g., Hesp and Shepherd, 1978; Heerdegen and Shepherd, 1992; Begg et

al., 2005; Litchfield and Berryman, 2005). Equivalent LGM aggradation surfaces

have been identified plunging beneath the Holocene fill in many other locations

throughout New Zealand (e.g., Suggate, 1968; Shepherd et al., 1986; Brown et al.,

1988; Brown, 1995; Bal, 1996; Dravid and Brown, 1997; Berryman et al., 2000;

Litchfield, 2003; Litchfield and Berryman, 2005, 2006; Begg et al., 2015).

In the north and east of the valley the LGM aggradation gravel unit is located

within 5-10 m of the surface of the floodplain (Fig. 3A-D, above 14 km along the

cross-sections), with the depth to the unit increasing to ~20-30 m beneath the

floodplain surface in the vicinity of Opiki (Fig. 3C-D, between 0-4 km along the

cross-sections). Downstream of Opiki the fluvial gravel facies is initially confined

to the centre and true left (southern side) of the valley at depths of 15-25 m

below PMSL (Fig. 4A, between 15-26 km along the cross-section, and Fig. 4B,

between 23-26 km along the cross-section). The gravel facies then switches to the

true right (northern side) of the valley and dips in a north easterly direction,

reaching depths of ~50-60 m below PMSL at the coast (Fig. 4B, between 0-14 km

along the cross-section, and Fig. 4C, between 0-12 km along the cross-section).

The geographic position of the fluvial gravels facies is interpreted to represent the

position of the primary channel belt of the Manawatu River during the LGM.

The fluvial gravels facies is interpreted to have been deposited between 30,000-

18,000 years BP, following the age constraints presented for LGM river

aggradation in the southern North Island (e.g., Litchfield, 2003; Litchfield and

Berryman, 2005; Clement and Fuller, 2007). Lieffering (1990) presented a

radiocarbon age of 29,780 ± 1600 cal. yr BP from a tree trunk (Tab. 1) recovered

~22 m beneath the floodplain surface from a silt and clay unit sandwiched

between two gravel units (Fig. 3B, at 10 km along the cross-section length). This

age falls at the onset of LGM aggradational river behaviour in the southern North

Island (e.g., Litchfield and Berryman, 2005; Clement and Fuller, 2007). The

overlying gravel unit is therefore inferred to be the LGM-age fluvial gravels facies,

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with the underlying gravel unit potentially gravels of MIS 3 age (e.g., Begg et al.,

2005; Clement and Fuller, 2007). The silt-clay unit underlying the fluvial gravels

facies is laterally extensive upstream of Opiki (Fig. 3), and is inferred to mark the

base of the fluvial gravels facies. Towards the coast silts and clays infrequently

underlie the LGM gravels, the base of which is therefore uncertain (Fig. 4). Hesp

and Shepherd (1978) reported a radiocarbon age of 41,500 ± 4,900 14C years

from a totara log (Tab. 1) embedded within a single 20 m thick gravel unit. The

upper part of this unit is interpreted to be the LGM fluvial gravels facies directly

overlying older MIS 3 fluvial deposits within which the totara log was embedded.

Organic-rich sands, silts, and clays (BHD)

Immediately downstream of Opiki a number of water-wells penetrate organic-rich

sand, silt, and clay units (Fig. 4A, above 20 km along the cross-section, and 3B,

above 23 km along the cross-section). These units are recorded by well-drillers

with descriptions such as ‘blue clay – timber and peat’, ‘grey silt and vege, some

decayed wood’, ‘soft blue silty clay, wood, and vege’, or simply ‘peat’. The peat,

wood, and vegetation content of these units is distinctive compared to the fluvial

sands, silts, and clays facies, which typically lacks description of organic material.

The organic-rich nature of this unit is interpreted to represent a low-energy

coastal wetland or bay-head delta environment proximal to the inflow of the

Manawatu River at the landward margin of the Holocene palaeo-estuary (e.g., Roy

et al., 1980, 1994; Chapman et al., 1982; Roy 1984, 1994; Roy and Boyd, 1996;

Nichol et al., 1997).

Fluvial clays and silts (TFL)

In the lower valley, within ~15 km of the present-day coastline, the LGM fluvial

gravels facies is overlain by ~1-7 m thick clay and silt units (Fig. 4B and C, between

0-14 km along the cross-sections). At the coast these units lie at elevations of ~50-

60 m below PMSL, rising to ~25 m below PMSL ~14 km up the valley. These units

are recorded by well drillers with descriptions such as ‘grey clays’, ‘blue clay’,

‘grey/brown clay’, and ‘blue silts’. Sand is infrequently reported as a co-

constituent, such as ‘sand and clay’ or ‘clay and sand layers’, as is organic

material. This facies is interpreted to represent retrogressive aggradation by the

Manawatu River during the early stages of the postglacial marine transgression

(e.g., Zhang and Li, 1996; Nichol et al., 1997; Li et al., 2002; Ishihara et al., 2012;

Bruno et al, 2017). As sea-level rose the elevated base level resulted in

adjustment of the river long profile, flattening the gradient in the lower valley,

reducing stream power and transport capacity, resulting in the deposition of the

fluvial silts and clays overlying the LGM fluvial gravels facies. Based on the

stratigraphic relationship with the underlying LGM fluvial gravel facies the fluvial

clays and silts facies is presumed to be younger than 18,000 years BP. Rapid sea-

level rise on the Manawatu coast associated with Meltwater Pulse 1A (e.g.,

Peltier, 2005; Gregoire et al., 2012) constrains the minimum age of the fluvial

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clays and silts at the mouth of the valley at c. 14,100 years, as this is

approximately when the mouth of the valley at ~60 m below MSL would have

been flooded by the marine transgression (Fig. 5). The RSL record also provides a

minimum age constraint for retrogressive aggradation by the Manawatu River at

the upstream limit of the facies at ~25-30m below MSL of c. 11,500 years BP, as

this is when this part of the valley would have been flooded by rising sea-levels

(Fig. 5).

Estuarine/marine sands, silts, and clays (EMS)

Sand, silt, and clay units with abundant shell, and occasionally wood or peat

content, are interpreted to have been deposited in the mouth and central basin of

the palaeo estuary. This facies is recorded in well logs with descriptions such as

‘fine silty/clayey sand with wood and shell fragments, minor clay seams’, ‘Sand,

odd shell and silt layer’, and ‘grey clay, shell, and sand’. The marine/estuarine

sands, silts, and clays facies encompasses a number of genetically-similar

sedimentary units such as transgressive sands, flood-tide delta sands, barrier and

nearshore sands, and estuarine central basin deposits (e.g., Roy et al., 1980;

Chapman et al., 1982; Roy, 1984a,b, 1994; Sloss el al., 2005, 2006a, 2010, 2007),

that cannot be reliably distinguished because of the paucity of detail in well

driller’s descriptions. The facies increases in thickness from ~15-20 m in the mid-

valley, to ~45-55 m thick at the coast (Fig. 4, between 0-22 km along the cross-

sections).

The presence of shells is taken as an indicator of a marine/estuarine depositional

setting, though not all sand, silt, and clay units in this facies contain shells. Shell

species are not generally reported by well drillers. Shepherd and Lees (1987)

drilled a borehole 4 km southwest of Opiki, penetrating 9 m of fluvial silts before

encountering a unit of at least 11 m thickness of thinly interbedded silty sand and

clays. This unit contained numerous A. stutchburyi shells (the only species

present) with no evidence of abrasion caused by transportation. A. stutchburyi is a

bivalve common to sheltered, estuarine locations, inhabiting the intertidal lower

foreshore (Gibb, 1979; Wilson et al., 2007; Clement et al., 2016). Shepherd and

Lees (1987) noted that the A. stutchburyi valves recovered from their borehole

were considered to have ‘lived in the extreme marginal low salinity for the

species, in other words, near the upper end of an estuary’ (A.G. Beu, pers. comm.

1980, reported in Shepherd and Lees, 1987). Shepherd and Lees (1987) reported a

radiocarbon age of 7090 ± 390 cal. yr BP (Tab. 1) from an A. stutchburyi valve

recovered near the top of the interbedded silt sand and clay unit, at ~5.6 m below

MSL.

At Foxton, near the mouth of the Manawatu valley, Te Punga (1958) presented a

radiocarbon age of 11,400 ± 230 cal. yr BP from a Podocarpus spp. log recovered

from ~45 m below MSL from a clay and sand unit directly overlying a unit of well-

rounded gravel (Fig. 4B, at approximately 4.5 km along the cross-section). The

relative sea-level (RSL) history for the Manawatu indicates that sea-level at this

time was ~30 m below PMSL (Clement et al., 2016), indicating that the

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Podocarpus log and clay and sand unit was deposited in a shallow

marine/estuarine environment with a water depth of up to 15 m. This provides a

minimum age for the establishment of marine/estuarine conditions at the mouth

of the valley and the marine/estuarine sands, silts, and clays facies.

Marine gravelly-sands (MGS)

Gravel and sand units frequently containing shells are recorded in well logs from

the mouth of the valley near Foxton (Fig. 4B and C, between 0-7 km along the

length of the cross-sections). These units dip towards the mouth of the valley, and

are described by well drillers as ‘blue gravels and sand’, ‘sand and gravel’, or ‘fine

blue sand and blue gravel’. The shell-content and large grainsize indicates a

higher-energy marine environment. These units are interpreted to represent a

high-energy wave-dominated beach/shoreface environment, similar to the

gravelly marine sand facies described by Shepherd et al. (1986) from the coast

near Tangimoana, approximately 15 km to the north of the Manawatu valley.

Shepherd et al. (1986) suggested an approximate maximum age of 5,000 years BP

for their gravelly marine sand facies, though the accuracy of their chronology is

unclear as many of their radiocarbon ages came from disarticulated valves within

the marine gravels, and so are likely to have been subject to reworking.

Fluvial sands, silts, and clays (FL)

The uppermost-units penetrated by water-wells throughout the Manawatu valley

are typically sands, silts, and clays (Figs 3 and 4). These units are variously

described by well drillers as ‘clay’, ‘silt’, or ‘sand’, or sometimes with more

descriptive lithologies such as ‘blue silty clay’, ‘grey/blue sandy silts’, or ‘silt and

sand layers’. These sands, silts, and clay units are typically free of wood, peat, and

other organic material. The facies varies in thickness from ~1-2 m in the upper

parts of the valley (Fig. 3, above ~16 km along all four cross-sections), and

increases in thickness in a down-valley direction to a maximum thickness of ~15-

20 m in the vicinity of Opiki (Fig. 3, between 0-4 km along all four cross-sections).

These sand, silt, and clay units are interpreted to represent the fluvial depositional

environment of the Holocene and modern floodplain. This interpretation is based

on the position of the water-wells within the gemorphologically- and geologically-

mapped extent of the floodplain (e.g., Fair, 1968; Hesp and Shepherd, 1978; Begg

and Johnston, 2000; Lee and Begg, 2002; Clement et al., 2010), and conforms with

descriptions of Holocene fluvial deposits in the Manawatu (e.g., Fair, 1968;

Heerdegen, 1972; Page and Heerdegen, 1985; Clement et al., 2010). In this facies

localised occurrences of gravel units are also preserved close to the surface of the

floodplain upstream of Opiki (Fig. 3B, between 0-5 km along the cross-section,

and 3C, between 11-14 km along the cross-section). Here these gravel units are

likely to represent the active channel belt of the contemporary Manawatu River

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floodplain, which has been characterised by lateral channel change in the past c.

2,000 cal. yr BP (LoRe et al., in review).

Dune sands (DS)

Close to the coast water-wells were drilled through the sands of the Foxton,

Motuiti, and Waitarere Holocene dune-building phases, with parabolic dunes

extending inland up to 20 km from the modern coastline (Fig. 4; e.g., Cowie, 1963;

Shepherd and Lees, 1987; Muckersie and Shepherd, 1993, 1995; Hawke and

McConchie, 2006). The surface sand units described in wells drilled within the

Holocene dunefield are interpreted to be Holocene dune sands. The sands are

recorded by well drillers with descriptions such as ‘brown/grey sands’, ‘brown

sand’, ‘grey fine sand’, or simply just ‘sand’. This makes it impossible to precisely

distinguish the base of the dune-sand unit from genetically-similar units that may

be underlying the dunes, such as beach sands. For this reason the base of the

dune-sand facies is represented as ‘uncertain’.

Vibracore facies

Cross-sections constructed using the vibracores, supplemented with details of

auger holes presented by Hesp (1975), were produced for the valley between

Opiki and Shannon (Figs. 6 and 7), and Shannon and Foxton (Figs. 6 and 8). Five

facies were identified in the vibracores and auger holes on the basis of lithological

description, sediment texture and colour, faunal assemblages, organic content,

contact relationships, and architecture. The characteristics of each of these facies

are summarised in Table 2.

Coarse estuarine/marine sands (FTD)

Coarse, clean sands were penetrated at the base of two vibracores on the true left

(southern side) of the lower valley between Koputaroa and Shannon,

approximately 10-12 km from the present-day coastline (Figs. 8 and 9). The coarse

sand retarded the vibracorer, with the unit penetrated to a maximum thickness of

only 0.3 m. Sediment analysis shows that the sands are medium- to coarse-

grained, indicating a moderate to higher energy environment, and moderately- to

poorly-sorted, indicating rapid deposition or a mixed energy regime (e.g., Heap

and Nichol, 1997; Roy et al., 2001; Sloss, 2005; Sloss et al., 2005, 2007, 2010, in

review). It is inferred that this sand unit represents a flood-tide delta

transgressing into the central basin of the palaeo estuary (e.g., Roy et al., 1980,

1994; Nichol, 1991; Nichol et al., 1997; Roy 1984, 1994). A single disarticulated A.

stutchburyi valve recovered from the top of the sand unit (Fig. 8) constrains the

minimum age of deposition to 5,730 ± 130 cal. yr BP (Tab. 1). This facies is

considered to be one of the constituent facies within the estuarine/marine sands,

silts, and clays facies (EMS) identified in the water well logs (Figs. 3 and 4).

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Estuarine intertidal sands and silts (ESS)

Fine sands interbedded with silts were present in almost all vibracores recovered

from the middle and lower valley (Figs. 7 and 8). Sedimentological analysis shows

that the sands and silts are both very poorly-sorted, suggesting deposition in a

moderate energy environment. Shells and shell hash are common in this facies

(Fig. 9). A. stutchburyi is the dominant fauna, typically found as disarticulated

valves, though articulated A. stutchburyi are not uncommon. The presence of A.

stutchburyi suggests a low-energy, intertidal environment (Gibb, 1979; Wilson et

al., 2007; Clement et al., 2016). Other fossil shells preserved in the interbedded

sand and silt unit include rare examples of the gastropod Amphibola crenata and

the bivalve Tellina spp.. Both inhabit intertidal mudflats and shallow-water

environments (Morton and Miller, 1968; Powell, 1979; Juniper, 1981; Pechenik et

al., 2003). Based on the sedimentological and faunal characteristics of this facies,

it is interpreted to represent an estuarine intertidal flat environment within the

central basin of the palaeo-estuary. Radiocarbon age determinations from

articulated and disarticulated A. stutchburyi preserved in the estuarine silts and

sands range from 8,300 ± 320 cal. yr BP to 4,700 ± 160 cal. yr BP (Tab. 1; Figs. 7

and 8). This facies is considered to one of the constituent facies within the

estuarine/marine sands, silts, and clays facies (EMS) identified in the water well

logs (Figs. 3 and 4).

Irregularly-mixed organic-rich sands and silts (FBDt)

This facies is distinguished by the irregular mixing of discrete patches of fine- to

medium-grained sands and silts (hereafter described as ‘mottling’; Fig. 9). The silts

are typically olive brown, the sands dark grey. Both sands and silts are very

poorly-sorted. Fossil shells are absent, while wood, roots, and other organic

material is present to common. The mottled nature and irregular mixing of the

sand and silt components suggests a mixed fluvial/tidal energy regime, and is

interpreted to represent the tidally-influenced front of a bay-head delta

prograding into the head of the palaeo-estuary. Radiocarbon age determinations

from A. stutchburyi valves within the underlying estuarine sands and silts indicate

that in the middle valley the fluvial bay-head delta deposits are younger than

6,500 cal. yr BP, while in the lower valley the deposits are younger than 6,000-

5,500 cal. yr BP (Tab. 1; Figs. 7 and 8). This facies is considered to correlate with

the organic-rich sands, silts, and clays facies (BHD) identified in the water well logs

(Figs. 3 and 4).

Organic-rich silts and clays (FBDf)

This organic-rich silt and clay unit is comprised of poorly-sorted, near

symmetrically-skewed silty clays or clayey silts. The unit is characterised by the

presence of wood, roots, and other organic content (Fig. 9). This unit is

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interpreted to represent the fluvially-influenced delta plain environment of the

fluvial bay-head delta prograding into the estuary. This facies is considered to

correlate with the organic-rich sands, silts, and clays facies (BHD) identified in the

water well logs (Figs. 3 and 4).

Fluvial silts and clays (FLsc)

The uppermost unit penetrated by the majority of vibracores and auger holes

comprises poorly-sorted, near symmetrically-skewed silty clays or clayey silts (Fig.

9), suggesting rapid deposition and/or deposition under mixed energy conditions

(cf. Sahu, 1964; Valia and Cameron, 1977; Martins, 2003). This facies is

characterised by the absence of shells, wood, and other organic matter, and is

interpreted to represent fluvial deposition on the Holocene floodplain of the

Manawatu River. This facies is considered to correlate with the fluvial sands, silts,

and clays facies (FL) identified in the water well logs (Figs. 3 and 4).

Discussion

Facies architecture and Holocene evolution of the Manawatu incised-valley

valley system

Last Glacial Maximum to late-Pleistocene fluvial deposits (30,000-14,100 cal. yr

BP)

During the Last Glacial Maximum the climate of the North Island was colder, drier,

and windier than present, with vegetation cover being predominantly scrub and

grassland (Newnham et al., 2013). Enhanced physical weathering resulted in an

increased sediment supply to the Manawatu River (Fair, 1968; Clement and Fuller,

2007). The river, lacking the stream power to transport this sediment, aggraded,

and likely had a braided planform (Fig. 10A; Fair, 1968; Heerdegen and Shepherd,

1992; Clement et al., 2010). In the valley upstream of Palmerston North City, this

aggradation unit is preserved as a now-largely-eroded fluvial terrace, comprising

gravel (FG facies) and mantled with loess (Fair, 1968; Heerdegen and Shepherd,

1992). The LGM aggradation terrace plunges beneath the Holocene floodplain

downstream of Palmerston North, and may be traced underneath the Holocene

valley fill out to the coast (Figs. 3 and 4). In sequence stratigraphic-terms the

upper-contact of the LGM-age fluvial gravels facies constitutes the transgressive

surface for the Holocene valley fill (e.g., Zaitlin et al., 1994).

The fluvial gravel aggradation surface is mantled by the fine-grained fluvial clays

and silts facies. This represents a switch in the behaviour of the Manawatu River

in response to changing climatic conditions and rising sea-level. Post LGM climatic

amelioration resulted in expanding vegetation cover throughout the catchment

and increased rainfall (e.g., Alloway et al., 2007; Barrell et al., 2013). In the

catchment upstream of Palmerston North City the resulting increased stream

power and reduced sediment supply incised the fluvial gravels facies. In the lower

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valley proximal to the present-day coast, rapidly rising sea-levels associated with

Meltwater Pulse 1A raised the base level of the river from 77-54 m below PMSL

between 14,500-14,000 years BP (Fig. 5; Clement et al., 2016). This reduced

stream power in the lower valley, initiating the retrogressive (back-stepping)

fluvial aggradation of clays and silts during the initial transgressive system tract

(Fig. 10A).

Late Pleistocene to early-Holocene transgressive deposits (14,100 – 7,100 cal. yr

BP)

The Holocene RSL reconstruction for the Manawatu coast indicates that the

mouth of the incised valley at approximately ~50-60 m below PMSL would have

been inundated by rising sea-levels c. 14,000 years BP, during the final stages of

Meltwater Pulse 1A (Fig. 9B; Clement et al., 2016). Rapid rates of sea-level rise

between 14,000-11,500 years BP (Fig. 5) coupled with high rates of fluvial

sediment delivery allowed rapid deposition of the back-stepping transgressive

fluvial silts and clays on top of the LGM gravel aggradation surface in the seaward

half of the palaeo-estuary (Fig. 10B). The valley was progressively flooded by rising

sea-levels, with the marine transgression culminating between 7,240-6,500 cal. yr

BP (Fig. 5). During the early stages of the marine transgression sediment stored on

the continental shelf was transported shoreward and into the mouth of the

Manawatu valley, and deposited on top of the basal fluvial facies (Fig. 10C). This is

consistent with studies of drowned river valley and barrier estuaries from

Australia (e.g., Roy et al., 1994, Roy, 1994, Sloss et al., 2005, 2006a, 2006b, 2007,

2010). The radiocarbon age determination from the totara logs in the water well

near Foxton presented by Te Punga (1958; Tab. 1; Fig. 4B) indicates that ~7-8

meters of transgressive marine sand had accumulated in the mouth of the palaeo-

estuary by c. 11,400 cal. yr BP. This sand unit would have formed a large

subaqueous tidal delta in the mouth of the estuary (Fig. 10C; e.g., Roy et al., 1980;

Chapman et al., 1982; Roy 1984a, b, 1994; Sloss et al., 2005, 2007). Subsequently

wave energy would have driven vertical accretion of this proto-barrier as sand,

delivered from the north along the coast via longshore drift and the shoreface via

cross-shore sediment exchange, accumulated in the mouth of the estuary. The

marine and near-shore sediment supply would have also fed the flood-tide delta

as it migrated landward with rising sea-levels. On the true left (southern side) of

the valley, near Shannon, a low-energy estuarine tidal flat environment was

established c. 8,300 cal. yr BP. As sea-levels continued to rise estuarine conditions

were established progressively further up the valley. The maximum extent of the

palaeo estuary approximately 22 km up-valley from the present-day coastline was

achieved between 7,400-6,700 cal. yr BP, coincident with the peak of the RSL

highstand on the Manawatu coast (Fig. 5; Clement et al., 2016). The environment

at the head of the estuary at this time was a lower-energy, fluvially-dominated

environment, a proto-bay-head delta, characterised by rich vegetation preserved

as peats, wood and logs, and other organic matter (Fig. 10C).

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Early- to mid-Holocene transgressive and regressive deposits (7,100 – 4,700 cal. yr

BP)

Following the culmination of the marine transgression a low-energy intertidal

estuarine environment was established in the lower Manawatu valley as a barrier

formed in the mouth of the estuary, near Foxton. As sediment accumulated in the

mouth of the estuary the depositional environment there transitioned from a

deeper-water tidally-dominated environment characterised by marine sands, to a

higher-energy, shallow-water, wave-dominated environment characterised by the

marine sandy gravel facies (Fig. 10D). Vertical accretion of the marine sandy gravel

would have driven a transition from a sub-aqueous to subaerial barrier. This

barrier restricted marine influences in the lower valley, which would otherwise

have been open and aligned to the westerly approach of open ocean wave energy

(Clement et al., 2010). Flood-tide delta sands in the lower valley point to the

maintenance of an inlet channel, with the flood-tide delta expanding into the

lower estuary as it was fed with marine sands delivered to the inlet-mouth by

longshore drift. The transition from coarse sands to interbedded silts and sand

recorded in the vibracores may represent growth of the barrier at the estuary

mouth, and formation of a more sheltered environment within the estuary (e.g.,

Allard et al., 2009; Raynal et al., 2010). This suggests that the barrier may have

achieved subaerial expression c. 5,600 cal. yr BP. Whole A. stutchburyi valves in

growth position in vibracores recovered between Opiki and Shannon, point to

low-energy estuarine conditions in the upper reaches palaeo estuary, supporting

long-held suggestions that a barrier in the vicinity of the Himitangi Anticline north

of Foxton protected the palaeo estuary from marine influences at the height of

height of postglacial marine transgression (e.g., Te Punga 1953; Rich, 1959; Hesp,

1975; Hesp and Shepherd, 1978; Heerdegen and Shepherd, 1992; Clement et al.,

2010; Clement, 2011). To the north and south of the palaeo estuary parabolic

dune-building was initiated as sea-levels stabilised, with the dunes initially formed

of transgressive sands transported across the continental shelf by rising sea-levels

and then moved onshore by the dominant westerly swell and north-westerly

winds (Clement et al., 2010). Dune-building was sustained by the longshore supply

of sand to the Manawatu coast.

Infilling at the head of the estuary began immediately following the onset of the

RSL highstand. Fluvial sediment delivery by the Manawatu River to the estuary

was extremely rapid, documented by the near-instantaneous transition from

estuarine intertidal sands and silts to fluvial silts and clays (Fig. 10E). By c. 6,000

cal. yr BP rapid infilling had advanced the head of the estuary down-valley by

approximately 10 km, placing it just to the west of Shannon. By c. 4,700 cal. yr BP

the estuary basin had completely infilled, with an intertidal estuarine environment

established in the position of the mouth of the palaeo-estuary in the vicinity of

the Whirokino Cut (Figs. 1 and 8B), approximately 4 km from the present day

coastline.

Mid- to late-Holocene regressive deposits (4.700 cal. yr BP – present)

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Following the end of the sea-level highstand c. 4,000 cal. yr BP sea-level fell to

PMSL. This fall in RSL coupled with the high marine sediment supply from

longshore drift along the coast fuelled rapid coastal progradation (Fig. 10F). In

addition, with the infilling of the estuary accommodation space fluvial sediment

bypassed the infilled-valley system, transporting fluvial sediment from the

Manawatu River directly to the coastal margin. As a result the apex of the marine

gravelly sand facies, interpreted to be subaerial ‘core’ of the barrier, occurs

approximately 4 km from the present day coastline. Seaward of the barrier core

lies a wedge of shelly marine sand. This is consistent with previous suggestions of

the amount of Holocene coastal progradation along the Manawatu coast (e.g.,

McFadgen, 1974; Heerdegen and Shepherd, 1992; Muckersie and Shepherd,

1995), as well as suggested amounts of progradation of the coast north at

Tangimoana (~10 km north of the Manawatu River mouth; Shepherd et al., 1986;

Heerdegen and Shepherd, 1992), and south of Levin (~12 km south of the

Manawatu River mouth; e.g., Te Punga 1962; Fleming, 1965, 1972; Gibb, 1978;

Bernett, 1984; Palmer et al., 1988; Hawke and McConchie, 2005, 2006). An age

range of 5,000-6,000 years BP for the subaerial expression of the barrier suggests

an average coastal progradation rate through this period of ~0.66-0.80 m per year

along the Manawatu coast. The sand supplied to the Manawatu coast by

longshore drift and cross-shore sediment exchange continued to supply the

parabolic dunefield, which ultimately transgressed approximately 19 km from the

present-day coastline (Shepherd and Lees, 1987; Muckersie and Shepherd, 1993,

1995; Hawke and McConchie, 2006; Clement et al., 2010).

Comparisons of the facies architecture, morphostratigraphy, and sedimentary

evolution of the lower Manawatu valley with tripartite facies models of incised-

valley estuary evolution

The sedimentary fill and Holocene geomorphic evolution of drowned incised-

valleys is complex, reflecting the interplay between fluvial, wave, and tidal

processes, geomorphic responses to relative sea-level changes, antecedent

geomorphology, neotectonics, and controls on sediment flux. Over the past 40

years the research utilising facies analysis and sequence stratigraphy has resulted

in a number of widely recognised tripartite facies models of estuarine systems

that broadly differentiate between wave- and tide-dominated end-member forms

(e.g., Dalrymple et al., 1992; Zailtin et al., 1994; Boyd et al., 2006; Sloss et al.,

2005, 2007). Following the model presented by Roy et al. (1980) the lower

Manawatu valley is analogous to a drowned river valley estuary: a deeply-incised

valley with a deep-water entrance. Under the model proposed by Dalrymple et al.

(1992), the Manawatu corresponds to a wave-dominated estuary (Fig. 11).

However, the Manawatu River estuary represents a mature stage (infilled) system

now dominated by sediment bypass.

Tripartite facies models of wave-dominated estuaries recognise three main

depositional environments with contrasting energy regimes (e.g., Roy et al., 1980,

1984; Dalrymple et al., 1992; Sloss et al., 2005, 2007). At the estuary mouth lies a

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marine-process-dominated zone in which wave and tidal energy is concentrated.

The mouth of shallower wave-dominated incised-valleys is typified by a barrier

complex consisting of a barrier spit, aeolian dunes, back-barrier sand flats,

washover deposits, flood- and ebb-tide deltas, and a tidal inlet channel (e.g., Roy

et al., 1980; Nichol, 1991a; Dalrymple et al., 1992; Reinson, 1992; Roy, 1994; Heap

et al., 2004; Sloss et al., 2005, 2006a, 2006b, 2007, 2010). In more deeply-incised

drowned valleys the estuary mouth is characterised by expansive subaqueous

tidal deltas, though more mature deeply-incised systems may evolve to a barrier

complex (e.g., Roy et al., 1980; Roy 1984, 1994). The central portion of wave-

dominated estuaries is occupied by a central depositional basin influenced by a

mix of marine and fluvial processes, and characterised by either fine-grained

organic-rich muds or intertidal flats. The landward portion of the estuary is

dominated by fluvial processes and characterised by a floodplain/bay-head delta

complex prograding into the estuarine central basin (e.g., Roy et al., 1980, 1994;

Chapman et al., 1982; Roy, 1984a,b, 1994; Roy and Boyd, 1996; Sloss et al., 2005,

2007).

In recent years studies of wave-dominated estuarine systems have sought to

understand and contextualise deviation or variation from the idealised tripartite

facies models (e.g., Heap and Nichol, 1997; Dillenburg et al., 2000; Cooper, 2001;

Lobo et al., 2003; Fitzgerald et al., 2005; Brothers et al., 2008 Burningham, 2007,

2008; Wilson et al., 2007; Abrahim et al., 2008; Kennedy et al., 2008; Chaumillon

et al., 2010; Peterson and Phipps, 2016; Gregoire et al., 2017). These studies

placed an emphasis on understanding the influence of and interplay between

system boundary conditions (antecedent geomorphology, marine and fluvial

processes, sediment regime, relative sea-level change, tectonics, and

accommodation space) as factors controlling the geomorphic and

sedimentological evolution of wave-dominated estuarine systems (e.g., Roy et al.,

1980, 1994; Roy 1984, 1994; Dalrymple et al., 1992; Heap and Nichol, 1997; Sloss

et al., 2005, 2006a, 2006b, 2007, 2010). As Wilson et al. (2007) noted, the case

studies of Holocene incised-valley infilling from which the tripartite facies models

were distilled typically come from estuaries of extensive size (40-100 km length)

on relatively tectonically-stable coastlines, with relatively low river discharges

and comparatively limited sediment supplies (e.g., Roy et al., 1980, 1994;

Dalrymple et al., 1992; Allen and Posamentier, 1993, 1994; Heap and Nichol,

1997; Sloss et al., 2005, 2006a, 2006b, 2007, 2010). In contrast, the Manawatu

valley is sited in a tectonically-active landscape, is subject to comparatively large

river discharges, and receives large sediment inputs from both fluvial and marine

sources (e.g., Clement et al., 2010).However, this is not a guarantee of rapid

infilling, as shown by the Columbia River estuary, which while it received large

sediment loads, maintains unfilled accommodation space as a consequence of

winter wind-wave stripping of sediment in the central basin, and sediment bypass

of fluvial sediments directly to the ocean (e.g., Peterson et al., 2013, 2014).

Dalrymple et al. (1992) suggested that the lowstand fluvial deposits at the base of

an incised-valley’s sedimentary fill will be directly overlain by back-stepping fluvial

delta sediments that transgress up-valley with rising sea-levels (Fig. 11; see also

Roy et al., 1980, 2001; Chapman et al., 1982; Roy 1984a, b, 1994; Nichol et al.,

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1994; Zaitlin et al., 1994). Deposits of this back-stepping fluvial delta facies are

typically preserved along the full length of incised-valley estuaries (Fig. 11A; e.g.,

Roy, 1994, Nichol et al., 1997), as it is normally only the upper portion of the

transgressive succession that is susceptible to shoreface erosion (Dalrymple et al.,

1992). In the Manawatu incised-valley system the fine-grained silt and clay units

overlying the fluvial gravels facies in well logs from the lower valley record this

phase of fluvial aggradation. However, in the Manawatu incised-valley system the

back-stepping fluvial facies is only preserved within the fill of the seaward half of

the palaeo-estuary, below ~25 m below PMSL (Fig. 11B). The absence of the fluvial

facies in the landward half of the palaeo-estuary, and a dominance of intertidal

deposits, may reflect an increase in the rate of sea-level rise between c. 11,500-

11,000 years BP (Fig. 5). During this period rapidly rising sea-levels may have

outpaced the rate of fluvial sediment supply as the shoreline transgressed

through this section of the incised-valley, resulting in poor preservation of the

facies. By comparison, sediment flux and the lower rate of sea-level rise between

14,000-11,500 years BP led to the deposition and preservation of the back-

stepping fluvial facies in the seaward portion of the palaeo estuary (e.g., Davis and

Clifton, 1987; Zaitlin, 1994).

In contrast to studies that identify a back-stepping fluvial unit deposited during a

transgressive system tract, Sloss et al. (2005, 2006a, b, 2007) identified a

transgressive marine sand-sheet as the initial transgressive unit overlying the late-

Pleistocene antecedent surface in shallow incised-valley systems on the southeast

coast of Australia. The transgressive sand-sheet was deposited inside these

incised-valley systems as rising post-glacial sea levels breached Last Interglacial

relict barriers at the valley mouths. On the Manawatu coast Clement et al. (2010)

noted that the major source of sediment for Holocene dune development is

commonly cited as transgressive marine sands moved onshore by rising sea levels

during the latter stages of the post-glacial marine transgression. Earlier studies in

the Manawatu did not record the presence of a sediment unit analogous to these

transgressive sands (e.g., Hesp, 1975; Hesp and Shepherd 1978; Shepherd and

Lees, 1987). However, these earlier studies did not have access to stratigraphic

data, such as the water well logs, that penetrated the full thickness of the

Holocene sedimentary succession in the lower valley. The water well logs from

the lower valley clearly show the ingress of transgressive marine sands,

distinguished from fluvial sediments by the presence of marine fauna (Fig. 4B and

C), into the valley mouth prior to the culmination of the marine transgression,

forming an expansive tidal delta (Fig. 11B). This is consistent with other studies of

drowned river valley evolution (e.g., Roy et al., 1980; Roy 1984, 1994; Nichol et al.,

1997). In the Manawatu it remains unclear whether that in addition to forming

the tidal delta, transgressive sands formed a laterally extensive sand-sheet of the

type described by Sloss et al. (2005, 2006a, b, 2007).

During the early stages of the marine transgression the environment at the mouth

of the valley is interpreted to have been influenced by tidal rather than wave

processes. This is similar to other drowned river valley estuaries (e.g., Roy, 1994;

Allard et al., 2009), and reflects the inherited character of the deeply-incised

valley mouth which was occupied by a large subaqueous tidal delta. This delta

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would have reduced incident wave energy penetrating into the central basin of

the estuary as waves shoaled over the delta, but would have had no effect on

reducing the tidal range within the middle and upper reaches of the estuary.

During the sea-level highstand the rapid accumulation of sediments in the mouth

of the estuary eventually resulted in the formation of a sand and gravel barrier.

The Manawatu is unique in this respect, compared to examples of drowned river

valley estuaries from Australia which typically have not achieved this mature stage

of development (e.g., Roy 1984, 1994).

As sea-levels stabilised during the RSL highstand and sediments accumulated in

the mouth of the valley to form the proto-barrier, wave processes would have

supplanted tidal processes at the valley mouth. Formation of the barrier at the

mouth of the estuary, c. 5,600 cal. yr BP, as suggested by the change in energy

conditions within the central basin recorded by the transition from coarse, clean

estuarine/marine sands to estuarine intertidal sands and silts, was comparatively

early within the Holocene infilling history of the estuary (e.g., Roy 1984, 1994;

Allard et al., 2009; Raynal, 2010), particularly when considering that the

Manawatu valley was comparatively deeply-incised. The estuarine tidal flat

environment that is characteristic of the central basin suggests that while the

barrier may have effectively eliminated wave energy into the central reaches of

the estuary, the basin remained tidally-influenced. This is analogous to the

contemporary river, which is tidal inland as far as Shannon.

The estuarine sand, silt, and clay facies identified in the Manawatu is broadly

similar to the estuarine mud facies detailed by Roy (1984a), the central basin mud

facies described by Dalrymple et al. (1992), and the tidal estuarine sand and mud

facies defined by Allen and Posamentier (1993, 1994). However the facies

described by Roy (1984a), Dalrymple et al. (1992), and Allen and Posamentier

(1993, 1994) are predominantly muds and silts laid down in a deep central basin

(Fig. 11A). In contrast the central basin mud facies in the Manawatu valley is

absent; sands are ubiquitous, and even occasionally gravels are present (Fig. 11B).

This variation is interpreted to represent the higher fluvial and marine

sedimentation rates within the central basin of the Manawatu incised-valley. The

river inflow and sediment regime is significant (maximum flows of 4,500 m3 s-1,

current sediment yield is 3.7 Mt yr-1; Hicks et al., 2011), as is the influx of sediment

from marine sources: for example, the flood-tide delta transgressed 11-12 km into

the palaeo estuary from the present-day coastline (Fig. 8). The majority of the

suspended sediment content from the river is likely being exported directly to the

continental shelf, a situation reminiscent of the Gironde Estuary (e.g., Castaing

and Allen, 1981; Lesueur et al., 1996).

Increasingly studies of incised-valley infilling from New Zealand demonstrate that

the development and preservation of the central basin facies is affected by the

combined effects of the allocentric controls of accommodation space,

neotectonics, relative sea-level change, and sedimentation rate (e.g., Heap and

Nichol, 1997; Wilson et al., 2007; Abrahim et al., 2008; Kennedy et al., 2008). In

Tamaki Estuary, Auckland, Abrahim et al. (2008) reported the absence of a central

basin facies in a sequence that is otherwise broadly consistent with the idealised

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facies model for a wave-dominated incised-valley system. This was attributed the

restricted form (~500 m wide) of the incised-valley. Wilson et al. (2007) also

documented poor preservation of central basin sediments in the similarly narrow

Pakarae River estuary (<500 m wide), on the tectonically active Raukamarua

Peninsula, on the North Island’s east coast. This poor preservation was attributed

to rapid fluvial progradation into Pakarae River estuary in direct response to

coseismic uplift events, leading to the truncation of central basin deposits in the

sedimentary sequence. Heap and Nichol (1997) working in Weiti River Estuary in

the North Island and Kennedy et al. (2008) working in Whanganui Inlet in the

South Island both reported high sedimentation rates in shallow incised valley

systems and the lack of a seaward barrier that limited the development of a low-

energy central basin mud facies.

In the Manawatu the deposition of estuarine sediment into the central basin was

extremely short lived (Figs. 7 and 8). The maximum extent of the palaeo-estuary

was achieved between 7,400-6,700 cal. yr BP, approximately 22 km up-valley from

the present-day coastline (Fig. 7). By 4,700 cal. yr BP the central basin had been

completely infilled, a process that took only 2,000-2,700 years (Fig. 8). Though the

lower Manawatu valley is not particularly narrow, or prone to coseismic uplift

events, it does possess an extremely high sediment load. Infilling by this sediment

load would have been aided by RSL regression following the peak of the highstand

c. 7,400 cal. yr BP. This is similar to the sedimentation histories described by Heap

and Nichol (1997) and Kennedy et al. (2008). Clement et al. (2010) noted that

both Hesp (1975) and Hesp and Shepherd (1978) documented an abrupt

transition from estuarine tidal flats (equivalent to the central basin facies) to

alluvium. However, neither described an intermediary bay-head delta sequence. It

was therefore hypothesised that the bay-head delta facies may be missing from

the lower Manawatu valley, perhaps due to reworking by fluvial processes.

Alternatively, Clement et al. (2010) suggested that rapid and high rates of

sediment delivery by the Manawatu River may have inhibited bay-head delta

formation, as rapid infill of the central basin eliminated the necessary

accommodation space, thereafter leading to sediment bypass. Results from this

research show that a bay-head delta did prograde into the central basin (Figs. 7

and 8); however it is identified only intermittently and appears to have been a

very short-lived feature in the rapid transition between estuarine tidal flats and

the floodplain environment (Fig. 11B). Whether this is a consequence of rapid

sedimentation and sediment bypass or the bay-head delta moving about the

valley as it prograded is unclear. However, this does not alter the conclusion

regarding the impact of rapid sediment delivery by the Manawatu River during

the mid to late Holocene aided by the fall in RSL during this period. The Manawatu

incised-valley estuary infilled extremely quickly, and the estuary's central basin

was a very short-lived feature, with a total elapsed time of infilling of only c.

2,000-2,700 years (Figs. 7 and 8). Wind-wave stripping of central basin sediments,

as experienced in other high sediment load estuaries (e.g., Peterson et al., 2013,

2014), does not appear to feature in the Manawatu, or the effect is overwhelmed

by the high sediment supply.

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The scenario of rapid sediment infilling is also consistent with post-LGM modelled

sediment flux identified elsewhere in New Zealand. Upton et al. (2013) identified

the period from 12,500-5,500 years BP as being the most productive period of

sediment generation in the Waipaoa catchment following the LGM, before

adjustment to higher sea-levels reduced stream power and sediment export from

the catchment, essentially due to infilling of the lower valley. This is analogous to

the situation in the Manawatu. The modelling by Upton et al. (2013) suggested

that the sediment load of the Waipaoa River increased with climate warming

following the Antarctic Cold Reversal (14,500-12,500 years BP; Wilmshurst et al.,

2007), because warmer, wetter conditions enhanced river erosion, relative to the

colder, drier LGM conditions. In addition to bioclimatic conditions favouring

incision, the role of tectonics in enhancing postglacial fluvial sediment flux in the

Manawatu should also be recognised (though it is currently beyond

quantification), as considerable incision (including into underlying Miocene

mudstone bedrock) has taken place during the postglacial period, which would

also have generated an as yet undetermined quantum of sediment.

Conclusions

Water well logs and vibracores recovered from the lower Manawatu valley record

the facies architecture, morphostratigraphy, and sedimentary evolution of this

drowned river incised-valley estuary, covering the period of the past c. 30,000

years. Analysis of the sedimentary fill of the lower Manawatu valley reveals a

number of unique features of its evolutionary history compared to established

conceptual models of the sedimentary and geomorphological evolution of wave-

dominated barrier estuaries. These variations are interpreted to primarily reflect

the rapid infilling of the palaeo-estuary driven by the high rates of sediment

supplied to the valley from both marine and fluvial sources. Results in this study

also indicate that the central basin is characterised by sands and silts, rather than

the fine-grained central basin mud facies typical of incised-valley estuaries with

more limited rates of sediment input. The transition from the sandy central basin

estuarine environment to the overlying floodplain environment was extremely

rapid. In some locations in the Manawatu valley the transitional fluvial bay-head

delta is missing completely, while in most locations where it is identified it present

only as a very thin deposit, signifying this rapid environmental transition. The

Manawatu system matured extremely rapidly: infilling of the estuarine central

basin took only 2,000-2,700 years. In contrast to less mature examples of

drowned river valley estuaries with lower rates of sediment flux, the mouth of the

Manawatu palaeo-estuary rapidly evolved from a tidally-influenced subaqueous

delta to a wave-dominated barrier estuary. This research provides significant

additional detail for established models of the evolution of incised-valley systems,

specifically for systems with high sedimentation rates that have undergone rapid-

infilling to quickly achieve a mature stage of estuary evolution.

Acknowledgements

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This research was undertaken as part of a PhD by AJHC, supported by a Top

Achievers’ Doctoral Scholarship, the Massey University Research Fund, the

Geoscience Society of New Zealand Wellman Research Award, and funding from

the School of People, Environment, and Planning, Massey University. Many thanks

to Hisham Zarour (Horizons Regional Council) for providing the well log data, and

John Dymond (Landcare Research) for providing a subset of the ECOSAT DEM.

Thank you to the many farmers and landowners who allowed access to their land

for vibracoring. Thank you to David Feek, Gigi Woods, John Appleby, Robert

Dykes, and Jane Richardson for assistance in the field. Many thanks to Pippa

Whitehouse (Durham University) for providing the GIA-modelled RSL predictions

for the Manawatu coast. This manuscript benefitted significantly from a review by

Mark Macklin and the constructive comments of two anonymous reviewers. This

paper is a contribution to the INQUA Commission on Coastal and Marine

Processes.

Supplementary data

Details of the well logs and vibracores, including all lithological data used in this

analysis, are included in a ZIP file which accompanies the electronic version of this

paper.

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List of Figures and Tables

Figure 1:

Geological/geomorphological map of the lower Manawatu valley at 1:300,000

scale, showing localities, major features, and previous estimates of the extent of

the Holocene palaeo-estuary. Base geological map data from Heron (2014).

Figure 2:

Map of the lower Manawatu valley showing the locations of cross-sections

constructed from water well logs (Figs. 3 and 4).

Figure 3:

Cross-sections for the Manawatu valley upstream of Opiki constructed from water

well logs.

Figure 4:

Cross-sections for the Manawatu valley downstream of Opiki constructed from

water well logs.

Figure 5:

(A) Palaeo sea-level index points for the southwest North Island, including the

lower Manawatu valley, together with the GIA-modelled predictions of RSL

change, sea-surface height (SSH), and solid Earth deformation (DEF) for the

Manawatu River mouth for the past 20,000 years BP (after Clement et al., 2016).

(B) Enlargement of the past 10,000 years BP showing palaeo sea-level index points

and GIA-model predictions.

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Figure 6:

Map of the lower Manawatu valley showing the locations of cross-sections

constructed from vibracores.

Figure 7:

Vibracores cross-sections for the Manawatu valley from Opiki to Shannon.

Figure 8:

Vibracores cross-sections for the Manawatu valley from Shannon to Foxton.

Figure 9:

Representative sections of each of the five major facies identified in the

vibracores recovered from the lower Manawatu valley. Each photograph is of a

~150 mm section of core, approximately 50 mm wide. From left to right: A) Facies

FTD – coarse, clean estuarine/marine sands, interpreted as a flood-tide

delta/washover deposit. B) ESS – estuarine intertidal sands and silts, interpreted

as a tidal flat environment within the estuarine central basin. C) FBDt – irregularly-

mixed organic-rich sands and silts, interpreted as a mixed-energy fluvial/tidal bay-

head delta environment. D) FBDf – organic-rich silts and clays, interpreted as a

fluvially-dominated bay-head delta environment. E) FLsc – fluvial silts and clays,

interpreted as the Holocene and modern river floodplain.

Figure 10:

Conceptual summary sequence of the facies evolution of the lower Manawatu

valley from the Last Glacial Maximum until present: (A) Last Glacial Maximum to

late Pleistocene lowstand aggradation gravels overlain by backstepping fluvial silts

and clays deposited during the early stages of sea-level rise; (B) Deposits of the

early-Holocene marine transgression; (C) Holocene sea-level highstand; (D) Mid-

Holocene barrier development and initial infilling of the central basin by fluvial

deposits; (E) Mid-Holocene estuary infilling; (F) Late Holocene sediment bypass

and coastal progradation.

Figure 11:

Comparison between the wave-dominated estuary facies model of Dalrymple et

al. (1992), and a model for the Manawatu estuary. (A) Model presented by

Dalrymple et al. (1992) for a wave dominated estuary, featuring a thick estuarine

central basin facies comprised of fine-grained sediments, underlain and overlain

by laterally-extensive fluvial deposits. (B) Model of the estuarine facies of the

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Manawatu valley. Backstepping fluvial deposits of the transgressive systems tract

pinch out mid-way up the length of the valley. The central basin space is filled

with course-grained marine sediments deposited during the transgressive phase;

highstand central basin deposits occupy only a fraction of the central basin

accommodation space. The overlying bay-head delta deposit is comparatively

thin, and in places non-existent.

Table 1:

Details of radiocarbon ages from wood and shell samples recovered from

vibracores, well logs, and auger holes drilled throughout the lower Manawatu

valley.

Table 2:

Distribution and characteristics of sedimentary facies of the Manawatu valley

Holocene infill sequence identified in water well logs, vibracores, and auger holes.

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Figure 1

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Figure 2

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Figure 3

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Figure 4

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Figure 5

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Figure 6

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Figure 7

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Figure 8

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Figure 9

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Figure 10

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Figure 11

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Table 1: Details of radiocarbon ages from wood and shell samples recovered from

vibracores, well logs, and auger holes drilled throughout the lower Manawatu

valley.

Source

Radiocarbon laboratory

code Material dated

Facies

(cf. Table

2)

Sample depth

(m, relativ

e to MSL)

Originally

reported age

(years BP)

Conventional

radiocarbon age (14C

years)

Sidereal

age range (2-σ,

cal. yr BP)

This study Wk22171 A. stutchburyi1

ESS -2.14 → 7150 ± 41 7850-8320

This study Wk22172 A. stutchburyi1

ESS -3.55 → 7390±42 7990-8630

This study Wk22173 A. stutchburyi1

ESS +0.15 → 5910±37 6500-7140

This study Wk22176 Wood FLsc +2.28 → 4974±35 5590-5730

This study Wk26357 A. stutchburyi2

ESS -2.01 → 5588±30 5880-6150

This study Wk26358 A. stutchburyi1

FTD -4.10 → 5343±30 5590-5860

This study Wk26359 A. stutchburyi1

ESS -1.48 → 6365±30 6700-6990

This study Wk26360 A. stutchburyi2

ESS -1.42 → 6260±30 6590-6870

This study Wk26361 A. stutchburyi2

ESS -1.72 → 6733±30 7160-7390

This study Wk26362 A. stutchburyi2

ESS -2.81 → 5920±30 6240-6470

This study Wk26363 A. stutchburyi2

ESS -3.68 → 5320±30 5580-5840

This study Wk26364 A. stutchburyi2

ESS -3.70 → 5661±30 5930-6200

This study Wk28289 A. stutchburyi2

ESS -2.84 → 5717±30 5990-6260

This study Wk28299 A. stutchburyi2

ESS -0.94 → 6386±32 6710-7020

Te Punga (1958)

NZ81 Podocarpus log

EMS -45.72 9,900 ± 150

9919±76 11170-11620

Hesp and Shepherd (1978)

NZ3085 A. stutchburyi

ESS +1.10 6,630 ± 70

6475±45 6810-7150

Hesp and NZ3938 Totara log FG -34.88 42,700 ± 41509±4895 36620-

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Shepherd (1978)

7650 500003

Shepherd and Lees (1987)

NZ5218 A. stutchburyi

ESS -5.60 6,280 ± 220

6613±181 6700-7480

Lieffering (1990)

NZ7693 Tree trunk FG -6.04 25,900 ± 410

25953±879 28180-31390

M.J. Shepherd unpublished

NZ7911 A. stutchburyi2

ESS +0.02 → 4525±43 4530-4860

1 Disarticulated 2 Articulated 3 Impinges on the end of the calibration data set. 1-σ age range 41,780-49170 cal. yr BP.

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Table 2: Distribution and characteristics of sedimentary facies of the Manawatu

valley Holocene infill sequence identified in water well logs, vibracores, and auger

holes.

Facies Distribution Sedimentary characteristics, faunal assemblages, and organic content

Interpretation of depositional environment

Chronology of deposition

Water w

ell lo

gs

Fluvial gravels (FG) Widespread in water wells throughout the valley; at depths of 10-30 m below the floodplain up-valley of Opiki, falling to depths of ~-60 m below MSL at the coast; 10-20 m thick

Poorly sorted sandy or silty gravel; free of shells; occasional organic material (e.g., tree logs)

Fluvial aggradation during the Last Glacial Maximum

30,000-15,000 years BP

Organic-rich sands, silts, and clays (BHD)

Restricted to the mid-valley, between Opiki and Shannon; units 10-15 m thick

Sands, silts and clays with high organic content (wood, peat, or ‘vegetation’)

Coastal wetland or bay-head delta

c. 7,100-6,000 cal. yr BP

Fluvial clays and silts (TFL)

Lower 13-14 km of the valley, restricted to the true-right; unit 2-5 m thick, at depths of -30 to -60 m below MSL

Predominantly blue, grey, or brown clay; occasional silt content

Retrogradational fluvial aggradation during sea-level transgression

14,000-12,500 years BP

Estuarine/marine sands, silts, and clays (EMS)

Widespread in the valley downstream of Opiki and Shannon; units 20-60 m thick

Predominantly sands and silts, occasionally interbedded with clay; frequent shells; occasional organic content (wood or peat)

Estuary central basin; tidal delta/inlet; beach/shoreface

12,500 years BP to present

Marine gravelly-sands (MGS)

Restricted to the mouth of the valley, within ~4-7 km of the coast; units ~5-20 m thick; dipping from 0 to -10 below MSL at their inland extent, to -30 to -50 m below MSL at the mouth of the valley

Poorly-sorted gravelly sand or sandy gravel; frequent shells or shell hash; occasional organic content (wood and peat)

Barrier/beach/shoreface

c. 6,000 years BP to present

Fluvial sands, silts, and clays (FL)

Widespread; upper ~5-10 of water wells drilled across the floodplain

Poorly-sorted sands, silts, and clays; free of shells and organic materials; infrequent isolated gravel units

Holocene and modern river floodplain

Upper floodplain (up-valley of the maximum palaeo estuary extent), 15,000 years BP to present; lower floodplain (down-valley of the maximum palaeo estuary extent): 7,100 cal. yr BP to present.

Dune sands (DS) Widespread at the coast; upper ~5-10 m of water wells drilled within the dunefield

Brown or grey fine sand; free of shells; occasional organic material

Holocene parabolic dunes

c. 7,100 cal. yr BP to present. V

ibraco

res

Coarse, clean, estuarine/marine sands (FTD)

Lower valley, west of Shannon; only the upper 10-20 cm of this unit penetrated

Medium- to coarse-grained, moderately- to poorly-sorted sand, free of shells and organics

Flood-tide delta/washover deposit

7,100-5,600 cal. yr BP

Estuarine intertidal sands and silts (ESS)

Widespread in the lower valley downstream of Opiki and Shannon; upper 1-5 m of this unit penetrated

Fine sand interbedded with silt, both poorly sorted; broken and whole A. stutchburyi extremely common, A. crenata and Tellina spp. common, shell hash extremely common; free of organics

Estuarine central basin tidal flats

7,100-4,500 cal. yr BP

Irregularly-mixed organic-rich sands and silts (FBDt)

Lower valley downstream of Opiki; units 0.2-1.5 m thick; units lie at depths of +1 to -2 m MSL, decreasing in elevation toward the valley mouth

Irregularly-mixed fine-to-medium sands and silts (‘mottled’); wood, roots, and other organic material not uncommon; free of shells

Fluvial bay-head delta; mixed fluvial/tidal energy

6,200-4,500 cal. yr BP

Organic-rich silts and clays (FBDf)

Lower valley between Opiki and Shannon; units 0.3-0.4 m thick; units lie at depths of +8 to – 2 m MSL, decreasing in elevation in

Poorly-sorted, near symmetrically skewed silty clay or clayey silt; wood, roots, and other organic material common; no shells.

Fluvial bay-head delta; fluvially-dominated

6,200-4,500 cal. yr BP

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the down-valley direction Fluvial silts and clays (FLsc)

Throughout the lower valley downstream of Opiki; units 3-8 m thick

Poorly-sorted, near symmetrically skewed silty clay or clayey silt; free of organic content and shells

Holocene and modern river floodplain

c. 7,100 cal. yr BP to present

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Highlights

High fluvial and marine sediment delivery drove rapid infilling of the

estuary

The Manawatu estuary was completely infilled within 2700 years

The estuary lacks central basin muds and a bay-head delta facies

The Manawatu contrasts with estuaries with more limited rates of

sediment input

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