Late Quaternary changes in Amazonian ecosystems and their ... · The current role of Amazonia in...

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Palaeogeography, Palaeoclimatology, P

Late Quaternary changes in Amazonian ecosystems and their

implications for global carbon cycling

Francis E. Maylea,*, David J. Beerlingb,1

aDepartment of Geography, University of Leicester, Leicester LE1 7RH, UKbDepartment of Animal and Plant Sciences, University of Sheffield, Sheffield S10 2TN, UK

Received 27 June 2001; accepted 5 July 2002

Abstract

The current role of Amazonia in the terrestrial carbon budget is the focus of intensive scientific interest, in large part due to

its potential to accelerate global warming. However, its role in mediating CO2 changes over millennial time-scales since the last

glacial maximum (LGM) has generally been overlooked and is the subject of speculation. Recent advances in our understanding

of the Late Quaternary history of Amazonian ecosystems offers an opportunity to make more informed inferences about Late

Quaternary changes in the magnitude of Amazon carbon storage than has hitherto been possible. Therefore, in this paper, we

reconstruct changes in the magnitude of Amazon carbon storage over the last 21,000 years (since the LGM) by reference to

recently published palaeohydrological and palaeoecological data and compare these data with results from simulations using a

process-based terrestrial ecosystem model for the Mid-Holocene and the LGM. Building on these results further, we interpret

changes in tropical forest biomass in the context of Late Quaternary polar ice-core records of atmospheric methane and carbon

dioxide concentrations. Palaeo-data and model simulations show that Amazonia was predominantly forested at the LGM,

although there is evidence for savanna expansion near the margins of the Basin and southern Amazonia may have been covered

by deciduous/semi-deciduous dry forests rather than evergreen rain forests. We estimate Amazon C storage at the LGM to be

only 135 Gt C (50% smaller than today), but find that its proportion of the entire terrestrial carbon store was almost twice that of

today. The model shows that between the LGM and the Mid-Holocene there is a significant increase in evergreen broad-leaf

forests at the expense of deciduous forests and a 67% increase in total Amazon C storage, attributable to rising temperatures and

atmospheric CO2 levels. Although our results indicate that the Amazon Basin was dominated by rain forests throughout the

Holocene, rain forest cover expanded in the Late Holocene (at the expense of savannas) and total Amazon carbon storage is

simulated to have risen by 22% between the Mid-Holocene (225 Gt C) and the present day (Pre-Industrial) (225 Gt C).

Comparison of these Amazon carbon fluxes with palaeo-data from other parts of the world suggests that, contrary to previous

hypotheses, the terrestrial biosphere acted as a net carbon sink throughout the Holocene, and that the observed CO2 rise from

0031-0182/$ - s

doi:10.1016/j.pa

* Correspon

Edinburgh EH8

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alaeoecology 214 (2004) 11–25

ee front matter D 2004 Elsevier B.V. All rights reserved.

laeo.2004.06.016

ding author. Present address: Institute of Geography, School of Geosciences, University of Edinburgh, Drummond Street,

9XP, UK. Tel.: +44 131 650 2552; fax: +44 131 650 2524.

esses: [email protected] (F.E. Mayle)8 [email protected] (D.J. Beerling).

114 276 0159.

F.E. Mayle, D.J. Beerling / Palaeogeography, Palaeoclimatology, Palaeoecology 214 (2004) 11–2512

260 to 285 ppmv between 8 and 1 ka BP (revealed by the Antarctic Taylor Dome ice-core record) may have been driven by

release of carbon from the oceans rather than land.

D 2004 Elsevier B.V. All rights reserved.

Keywords: Amazon; Late Quaternary and Holocene; Carbon cycle; Vegetation-climate model; Palaeoecology; Carbon dioxide; Methane

1. Introduction

Amazonian tropical forests are estimated to

account for ~10% of the world’s terrestrial primary

productivity and 10% of the carbon stored in

terrestrial ecosystems (Melillo et al., 1993). They

therefore constitute a significant terrestrial carbon

reservoir or pool, which undoubtedly plays an

important role in the global carbon cycle. Tian et al.

(2000) estimate that the total carbon storage within the

undisturbed ecosystems of the Amazon Basin in 1980

was 127.6 Pg C (1 Pg=1015 g), of which 83%

occurred in tropical evergreen forests, 3.5% occurred

in tropical deciduous forests, and 4.5% occurred in

savannas. Different tropical ecosystems differ mark-

edly in terms of their net carbon storage per hectare

(i.e. vegetation+soils+litter). For example, Adams and

Faure (1998) obtained estimates of 320 tons C ha�1

for tropical rain forest vs. 260 tons C ha�1 for

seasonal dry forest (monsoon forest) vs. 90 tons C

ha�1 for savanna. Given that even short-term, inter-

annual fluxes in precipitation and temperature (e.g.

associated with El Nino events), and disturbance

regime (e.g. deforestation, fire) are believed to

significantly affect the role of Amazonian ecosystems

in the global carbon budget today (i.e. whether or not

Amazonia behaves as a net carbon source or sink)

(e.g. Phillips et al., 1998; Tian et al., 1998, 2000;

Houghton et al., 2000), it would be expected that any

significant changes in the geographic extent of

Amazonia’s ecosystems (e.g. rainforest vs. semi-

deciduous forest vs. savanna) over past millennia

would have markedly altered the size of the carbon

reservoir within Amazonia, and consequently have

had an important impact on the global carbon budget.

High-resolution ice-core records indicate that Late

Quaternary changes in Amazonian forest dynamics

were accompanied by marked shifts in atmospheric

concentrations of CO2 (Indermuhle et al., 1999) and

methane (Blunier et al., 1995) during the last

deglaciation and the Holocene. Reconciling these

millennial-scale fluxes in atmospheric carbon with

available evidence for changes in the terrestrial and

oceanic carbon budgets remains a considerable

challenge for carbon cycle modellers. The prospects

of meeting this challenge are, however, increasing as

we obtain a better quantitative understanding of the

influence of palaeoenvironments on carbon storage in

terrestrial ecosystems (vegetation and soils) (Beerling,

1999, 2000) and the effects of tropical forests on

climate (Crowley and Baum, 1997; Betts, 1999; Levis

et al., 1999; Kleidon and Stephan, 2001) and the

concentration of atmospheric CO2 itself (e.g. Cox et

al., 2000).

An example of this progress is provided by the

work of Indermuhle et al. (1999), who hypothesised a

cumulative release from the terrestrial biosphere of

195 Gt C between 7 and 1 cal ka BP (thousand

calendar years before present), of which 145 Gt C

were modelled to be taken up by the oceans and 50 Gt

C were taken up by the atmosphere (equivalent to the

25 ppm rise in atmospheric CO2 observed by

Indermqhle et al. in the Taylor Dome ice-core record

over this time). However, Beerling (2000) analyzed

this hypothesis using three independent lines of

evidence (global data-based reconstructions of terres-

trial carbon reservoirs, vegetation-climate modelling,

and high latitude stable isotope records) and showed

that neither approach supports a sufficiently large

terrestrial source of C to account for the observed rise

in atmospheric CO2 through the Holocene. Causal

mechanisms underlying Holocene atmospheric meth-

ane fluxes have also not been resolved, with

uncertainty over the relative roles of high latitude

and tropical wetlands (Chappellaz et al., 1993;

Ruddiman and Thomson, 2001).

The current role of Amazonia in the terrestrial C

budget is clearly the focus of great scientific interest,

given its potential to accelerate global warming (Cox

et al., 2000), which in turn depends on the response of

F.E. Mayle, D.J. Beerling / Palaeogeography, Palaeoclimatology, Palaeoecology 214 (2004) 11–25 13

the tropical forests to current CO2 and climate change

(Mahli and Grace, 2000). However, its role in

mediating CO2 changes over millennial time-scales

since the last glacial maximum (LGM, ca. 21 cal ka

BP) has generally been overlooked, in part perhaps

due to the scarcity of palaeoecological data from this

vast area. Recent advances in our understanding of the

Late Quaternary history of Amazonian ecosystems,

especially with respect to forest–savanna dynamics

(e.g. Behling and Hooghiemstra, 2000; Mayle et al.,

2000; De Freitas et al., 2001) offers an opportunity to

make better informed inferences about Late Quater-

nary changes in the magnitude of Amazon carbon

storage than has hitherto been possible.

In this context therefore, the aim of this paper is to:

(i) reconstruct changes in the magnitude of Amazon

carbon storage over the last 21000 years (since the

LGM) by reference to recently published palaeohy-

drological and palaeoecological data, (ii) use these

data for comparison with simulations using a process-

based terrestrial ecosystem model for the mid-Hol-

ocene and the last glacial maximum (Beerling, 1999),

(iii) interpret the palaeo-data and model evidence for

changes in biomass in the context of Late Quaternary

polar ice-core records of atmospheric methane and

carbon dioxide concentrations, and (iv) draw infer-

ences about the likely role of Amazonian ecosystems

in global carbon cycling since the LGM.

2. Palaeoecological evidence for fluxes in the

magnitude of the Amazon carbon reservoir

2.1. The last glacial maximum (ca. 21 cal ka BP, 1814C ka BP)

Although there are very few palaeoecological

records extending to the LGM, there is accumulating

evidence to indicate that most of the Amazon Basin

was forested at this time (Fig. 1), contrary to Haffer’s

(1969) hypothesis of widespread savanna with iso-

lated, disjunct forest refugia. Colinvaux et al. (1996)

and Bush et al. (2002) have provided pollen, sed-

imentological, and geochemical data that show that the

Lake Pata catchment in the central Amazon (0816VN,66841VW) was continuously forested over the last

170,000 years. De Freitas et al. (2001) collected soil

carbon isotope data from a 200 km transect, spanning

small, isolated pockets (ca. 10–100 km2) of seasonally

flooded savannas, between Porto Velho (Rondonia

State) and Humaita (Amazonas State), Brazil, between

the coordinates 8843VS, 63858VW and 7838VS,63804VW. These savanna dislandsT are surrounded by

rainforest and located ca. 450 km north of the southern

limit of rainforest communities in Bolivia. De Freitas

et al. show that these savanna islands at 17 14C ka BP

(ca. 20.5 cal ka BP) were no larger than today,

indicating that there was no expansion of savanna at

the expense of forest at this time.

However, in contrast to these sites in central and

southern Amazonia, records from sites closer to the

margins of the Basin do reveal changes in forest/

savanna distribution between the LGM and today.

Mayle et al. (2000) obtained a continuous 40,000 year

pollen record of vegetation change from Laguna

Chaplin (14828VS, 61804VW), Noel Kempff Mercado

National Park (NKMNP), and showed that Amazo-

nian rainforest communities during the LGM were

located at least 30 km north of their current southern

limit in eastern Bolivia, and that this ecotonal area was

then dominated by open savannas with rainforest

species restricted to riverine gallery forests. Pollen

evidence from Laguna El Pinal, in the Colombian

savannas of the Llanos Orientales, at the northern

margin of the basin (Behling and Hooghiemstra,

1999) signifies a virtually tree-less grassland sur-

rounding the lake at the LGM.

Whilst it would be unwarranted to extrapolate the

significance of these isolated data to Amazonia as a

whole, pollen data from cores of the Amazon Fan can

be considered a more reliable indicator of Basin-wide

changes in vegetation, since these sediments have been

deposited from the entire Amazon river catchment.

Pollen spectra from these Amazon Fan cores show no

significant changes in the relative proportions of forest

versus savanna pollen taxa between the LGM and the

Holocene (Haberle, 1997; Hoorn, 1997; Haberle and

Maslin, 1999). These findings corroborate the isolated

terrestrial pollen records, showing that while there is

evidence for more widespread savannas at the northern

and southern Amazonian margins relative to today

(e.g. eastern Bolivia, Mayle et al., 2000), most of the

Basin remained forested at the LGM.

However, determining the kinds of forests that

dominated Amazonia at the LGM is far from

straightforward, since most pollen types cannot be

Fig. 1. Map showing the location of sites discussed in the text. The shaded area shows the current distribution of Amazonian evergreen broad-

leaf forest (rain forest). The diagonal line between the coordinates 178S, 63VW and 38S, 45VW shows the approximate southern limit of

Amazonian rain forest, which we have used to define the southern limit of the Amazon Basin when modelling Amazon C storage and NPP. The

hatched area represents the Andes. dCal ka BPT refers to dthousand calendar years before presentT, where dpresentT refers to 1950 AD. dNKMNPTrefers to dNoel Kempff Mercado National ParkT.


identified beyond the genus level and many Ama-

zonian genera contain species from several different

ecosystems (e.g. rainforests, semi-deciduous forests,

savannas). Furthermore, since all forest species are C3

plants, bulk sediment stable carbon isotope analysis

cannot distinguish different types of forest community

(e.g. rainforests vs. dry forests). By analysing biogeo-

graphic patterns of disjunct tropical seasonally dry

forests (deciduous and semi-deciduous forests) in

South America, Pennington et al. (2000) raise the

possibility that dry forests, rather than rainforests,

dominated much of Amazonia during full glacial

times. Andean taxa, such as Podocarpus, are a

consistent feature of LGM lowland Amazon pollen

records, showing that Andean forest species invaded

much of the Amazon Basin, suggesting that this area

was significantly cooler than today, an interpretation

supported by analysis of noble gas concentrations in

fossil ground water in eastern Brazil, which show a 58C cooling compared to the present (Stute et al., 1995),

and tropical Atlantic sea-surface temperature (SST)

reconstructions (Guilderson et al., 1994).


The level of precipitation in Amazonia at the LGM

remains a contentious issue and is itself a feature

partly dependent on rainforests, which can recycle up

to 50% of their canopy transpiration as precipitation

(Shukla and Minz, 1982). If savannas and seasonally

dry forests were more abundant in Amazonia at the

LGM than today, this could be construed as evidence

of increased aridity, possibly attributable to reduced

evaporation from the tropical Atlantic Ocean and

reduced evapo-transpiration from the Amazon Basin

under cooler glacial conditions.

However, Baker et al. (2001a,b) show that Lake

Titicaca was overflowing (Fig. 1) and the Salar de

Uyuni salt flats were flooded at the LGM, indicating

that precipitation on the Altiplano of the Bolivian

Andes was even higher than today. Baker et al. assume

that because the Bolivian Andes receive most of their

precipitation from Amazonia today, the Amazon Basin

must therefore have been at least as wet at the LGM as

it is today. Such a scenario might at first glance appear

to contradict the available (albeit limited) palaeoeco-

logical data, which point to increased extent of

savannas at the Amazon margins at the expense of

rainforest, suggesting lower precipitation than present.

It might be argued that, unlike today, the climate

systems over the Bolivian Andes and Amazonia may

have been de-coupled during glacial times, with the

tropical Andes receiving predominantly winter precip-

itation from cold polar fronts originating from the

South Pacific Anticyclone over Patagonia, rather than

summer precipitation from Amazonia.

Alternatively, however, the lowland Amazon palae-

oecological recordsmay indeed tie in with Baker et al.’s

Andean precipitation record. Amazonian expansion of

deciduous forests and savannas at the expense of rain

forests at the LGM may not necessarily have been

caused by a reduction in precipitation, but instead by

lowered atmospheric CO2 levels (ca. 200 ppm), which

would be expected to have selectively favoured plants

using the C4 (e.g. savanna grasses) and CAM (e.g.

bromeliads, cacti) photosynthetic pathways over plants

(all trees) using the C3 pathway (Cowling and Sykes,

1999; Bennett and Willis, 2000), by virtue of their

greater photosynthetic efficiency, and hence water-use

efficiency (WUE) at low CO2 levels. Experimental data

have shown that C3 leaf and plant WUE is significantly

reduced in carbon-depleted atmospheres, in some

instances by up to 50–60% (e.g. Polley et al., 1993;

Cowling and Sage, 1998), due to higher rates of

transpiration as a result of increased stomatal con-

ductance. Low CO2 levels at the LGM (rather than

reduced precipitation) may therefore have been the

primary reason for expansion of savannas at the

ecotonal margins of the Basin and supports Penning-

ton’s hypothesis (2000) for replacement of evergreen

forest by deciduous and/or semi-deciduous forest.

Notwithstanding these uncertainties over Amazo-

nian precipitation and forest composition/structure, the

available palaeo-data clearly demonstrate that Ama-

zonia was predominantly forested at the LGM,

although there was greater extent of savannas at the

Amazon margins and there may have been more dry

forests (semi-deciduous or deciduous) and lower

canopy densities relative to today. This implies that

the Amazon C sink would have been reduced

compared to the present, although there are insufficient

data to quantify by how much.

2.2. Last glacial/Holocene transition (ca. 20–10 cal

ka BP, 17–9 14C ka BP)

There was considerable geographic variation in the

nature and timing of ecosystem changes across the

Amazon Basin over this interval. The pollen record

fromLaguna Chaplin, located at the rainforest–savanna

ecotone of eastern Bolivia (Mayle et al., 2000), shows

that savannas still dominated the landscape, but that

gallery forests now contained Podocarpus (a predom-

inantly Andean genus), and Alchornea (in addition to

Moraceae), suggesting reduced water stress (either due

to increased precipitation and/or WUE) and possibly

lower temperatures than before. Soil stable carbon

isotope data from a 200-km transect between Porto

Velho and Humaita, southern Amazonia (De Freitas et

al., 2001), show that this area had sufficient moisture to

support forest throughout this period. Ledru et al.

(2001) and Behling (2001) have shown that a Late-

glacial Podocarpus signal is common to many pollen

records, not just from southern Amazonia, but through-

out the Amazon Basin. It should be noted, however,

that there is considerable variability in the chronology

of this peak during this interval. For example, the

Podocarpus peak spans the LGM to the onset of the

Holocene in southern Amazonia (L. Chaplin, Mayle et

al., 2000; Carajas, Absy et al., 1991), the LGM to 14.214C ka BP in central Amazonia (L. Pata, Colinvaux et


al., 1996), and ca. 13.0–12.8 14C ka BP in eastern

Amazonia (L. do Caco, Ledru et al., 2001). In the

Colombian savannas of the Llanos Orientales border-

ing northern Amazonia, pollen data indicate that

climates did not become wetter until ca. 11.5 14C ka

BP or later (Behling and Hooghiemstra, 1998; 1999;

2001). Maslin and Burns (2000) used oxygen isotope

analysis of planktonic foraminifera from Amazon Fan

sediments to show that Amazon River discharge during

the Younger Dryas chronozone (11–10 14C ka BP) was

ca. 60% lower than today, implying an arid Amazon

Basin at this time. It is interesting that this episode of

apparent aridity correlates with the onset of wetter

conditions in the Colombian savannas and maximum

water levels in Lake Titicaca (Baker et al., 2001b),

which is difficult to reconcile.

There clearly appears to have been marked spatial

and temporal variability in the pattern of ecosystem

changes reflecting complex climatic evolution during

this glacial–interglacial transition. However, the

palaeoecological data, considered as a whole, indicate

that the Amazon Basin was more forested during this

interval than at the LGM, due to increased precip-

itation and/or increased WUE as CO2 levels rose. This

would suggest an increase in Amazon biomass and

hence C storage.

2.3. Early-mid Holocene

Stable carbon isotope data from soil organic matter

show expansion of savanna islands at the border of

Amazonas and Rondonia states, Brazil (Pessenda et

al., 1998a,b; De Freitas et al., 2001). In contrast to the

LGM, this savanna expansion is clear evidence for

early-mid Holocene aridity in southern Amazonia, as

CO2 levels were by now at least 60 ppm above LGM

levels and therefore no longer limiting for C3 plants,

and there is corroborating charcoal evidence for a

peak in fire frequencies between 7 and 3 14C ka BP in

Para State (eastern Brazilian Amazonia) (Turcq et al.,

1998; Soubies, 1979–1980) and pollen and charcoal

evidence for savannas and high fire frequencies at the

rainforest–savanna ecotone of eastern Bolivia

(NKMNP, Mayle et al., 2000). Further supporting

evidence for regional climatic aridity in southern

Amazonia through this interval comes from sediment

cores from Lake Titicaca in the Bolivian Altiplano,

which show that over the last 25,000 years the period

of lowest water levels and maximum aridity occurred

from 8.5 to 4.5 cal ka BP (Baker et al., 2001b).

Biogeographic shifts in the forest–savanna boundaries

of northern Amazonia appear to have been roughly in

phase with those of southern Amazonia (Behling and

Hooghiemstra, 1999, 2000, 2001), showing increased

savannas in the early-mid Holocene (until ca. 7–6 cal

ka BP) relative to the late-glacial period, although

higher lake levels at Laguna El Pinal and Laguna

Carimagua point to higher rainfall at the northern

Amazon margin than during the LGM.

Fossil pollen, charcoal, stable carbon isotope, and

lake level data all point to a basin-wide reduction in

precipitation in Amazonia, which caused replacement

of forest by savannas at the forest–savanna ecotones

and increased fires. Disappearance of Andean taxa (e.g.

Podocarpus) from the lowlands (reflected in all

Amazonian pollen records) also indicates that this

increase in aridity coincided with increased temper-

atures. This expansion of savannas at the expense of

forests suggests that Amazonian biomass (and hence

the C reservoir) was lower during the early-mid Holo-

cene than during the preceding Late-glacial period.

2.4. Late Holocene

Pollen data from Laguna Bella Vista and Laguna

Chaplin show that rainforest communities expanded

southwards to replace savannas in NKMNP, eastern

Bolivia, within the last three millennia to reach their

current geographical limit at ca. 158S (Mayle et al.,

2000). Furthermore, these data show that the present-

day rainforest boundary in eastern Bolivia constitutes

the southern-most extent of Amazonian rain forest in

South America over at least the past 50,000 years.

Corroborative evidence for forest expansion at the

expense of savannas at ca. 3 ka BP comes from stable

carbon isotope data from the Brazilian states of

Rondonia and Amazonas (De Freitas et al., 2001) and

pollen (Absy et al., 1991) and charcoal (Turcq et al.,

1998) data from Carajas, Para state, to the east. These

data clearly point to increased precipitation in southern

Amazonia within the last 3 millennia, also reflected in

rising water levels of Lake Titicaca (Cross et al., 2000;

Baker et al., 2001b) and increased precipitation on

Sajama Mountain (Thompson et al., 1998). This

encroachment of forest into savannas also occurred at

the northern margin of Amazonia, although this


expansion began somewhat earlier at ca. 7–6 cal ka BP

(Behling and Hooghiemstra, 1999, 2000, 2001). This

late Holocene forest expansion throughout the Amazon

Basin indicates that the Amazonian C store increased to

its greatest size over at least the past 21,000 years.

3. Integrating palaeodata and carbon storage

model estimates

The recent increase in the quantity and quality of

Amazonian palaeoecological data provides an oppor-

tunity to compare these observations in a qualitative

manner with patterns of Amazon vegetation distribu-

tion and C storage using process-based models of the

terrestrial carbon cycle (e.g. Cramer et al., 2001). Here,

we make this comparison using data extracted from

global equilibrium simulations reported for the Mid-

Holocene and the LGM (Beerling, 1999) using the

University of Sheffield Dynamic Global Vegetation

Model (SDGVM) (Woodward et al., 1995). This

SDGVM calculates vegetation properties under

steady-state conditions of climate and CO2, and

represents the physiological processes of plant nutrient

uptake, C3 and C4 photosynthesis, respiration, and

stomatal control of canopy transpiration. Above-

ground productivity is fully coupled to a below ground

model of soil carbon and nitrogen dynamics so that

plant litter (leaves and surface roots) is subsumed

through decomposition/nutrient cycling (Woodward et

al., 1998). The model predicts the distribution of

different plant functional types on the basis of annual

net primary production and biomass, competition for

light and other resources, as well as probability of

disturbance, and succession following disturbance

(Cramer et al., 2001).

In this analysis, we focused on three time intervals:

the Pre-Industrial, the Mid-Holocene, and the LGM.

The Pre-Industrial simulations of potential vegetation

in the Amazon Basin were made by forcing the

SDGVM with the historical land surface climatologies

produced by New et al. (1999, 2000) and averaging the

results for the period 1901–1910 AD. For each

palaeoclimate simulation, results are reported for two

GCM-derived climates, one with a high spatial

resolution (2.88 lat.�2.88 long.), the U.K. UniversitiesGlobal Atmospheric Modelling Programme

(UGAMP) GCM (Hall et al., 1996a,b; Hall and Valdes,

1997), and the other with a lower spatial resolution

(4.48 lat.�7.58 long.), the National Center for Atmos-

pheric Research (NCAR) community climate model

GCM (Kutzbach et al., 1998). The UGAMP GCM is

based on the European Centre for Medium-Range

Weather Forecasting model (ECMWF) and uses sea-

surface temperatures prescribed from palaeo-data. In

each case, model bias was minimized by calculating

locally differing monthly climate anomalies (i.e. GCM

control (present-day) run minus GCM palaeo-run) and

imposing these onto a modern underlying climatology

(cf. Beerling, 1999). Atmospheric CO2 concentrations

for the Pre-Industrial, Mid-Holocene and LGM were

300, 280 and 180 ppm, respectively.

Fig. 2 and Fig. 3 show the predicted distribution of

plant functional types and vegetation productivity and

land surface carbon storage in the Amazon Basin for

the (a) present-day (Pre-Industrial), (b) Mid-Holocene

(6 cal ka BP), and (c) LGM (21 cal ka BP). The Pre-

Industrial simulation is for the potential vegetation,

and does not account for anthropogenic land-cover

change (e.g. crops, urban areas, forest clearance). For

the purpose of calculating changes in Amazon C

storage simulated by the model (Table 1 and Table 2),

we do not refer to dAmazon BasinT sensu stricto (as

defined by the precise limit of its river catchment), but

instead refer to it in a more general sense as north of

the diagonal line between Santa Cruz, Bolivia (178S,63VW) and Sao Luis, Brazil (38S, 45VW) (Fig. 1),

which roughly corresponds to the present-day southern

limit of Amazonian evergreen broad-leaf forest. The

rationale for this approach is that we are primarily

interested in the land area presently covered by tropical

rain forest in South America (most of which happens

to occur within the Amazon drainage Basin, but also

extends into the highlands of, for example, Guyana

and Suriname) rather than the ecosystems that happen

to occur within the strict confines of the Amazon Basin

sensu stricto.

A key result from this series of equilibrium

ecosystem model simulations is that the Amazon

Basin remained predominantly forested over the last

21,000 years (Fig. 2a–c), which is supported by the

available pollen data (e.g. The Amazon Fan record).

Furthermore, deciduous broad-leaf forests (rather than

evergreen broad-leaf forests) are simulated to have

covered the southern half of the Basin at the LGM

(Fig. 2c), which supports the hypothesis of Pennington

Fig. 2. SDGVM model simulations forced with the UGAMP GCM. See text for full explanation. Key: c3, C3 grasses; c4, C4 grasses; ebl,

evergreen broad-leaf forest; enl, evergreen needle-leaf forest; dbl, deciduous broad-leaf forest; dnl, deciduous needle-leaf forest.

Fig. 3. SDGVM model simulations forced with the UGAMP GCM. See text for full explanation.


Table 1

Carbon storage changes in Amazonian vegetation biomass and soil organic matter since the last glacial maximum

Vegetation type Time interval

Pre-industrial

(1901–1910 average)

Mid-Holocene (1) Mid-Holocene (2) Last glacial

maximum (1)

Last glacial

maximum (2)

Vegetation Soils Vegetation Soils Vegetation Soils Vegetation Soils Vegetation Soils

Evergreen broad-leaved

forests

187.9 76.4 136.6 61.2 151.2 70.7 65.1 37.9 50.1 29.6

Deciduous broad-leaved

forests

4.6 3.9 11.5 8.1 0.0 0.0 19.9 14.8 26.1 20.9

C4 grasslands 0.2 0.6 0.1 0.4 0.0 0.0 0.05 0.4 0.2 1.9

C3 grasslands 0.0 0.0 0.1 4.5 0.1 5.1 0.07 1.9 0.0 0.0

Total 192.7 80.9 148.3 74.2 151.3 75.8 85.1 55.0 76.4 52.4

All figures are Gt C. The mid-Holocene and last glacial maximum climates (1) and (2) correspond to results from forcing the vegetation model

with climate simulated by the UGAMP and NCAR general circulation models, respectively.


et al. (2000), who used modern biogeographic data

from disjunct ecosystem distributions to postulate that

much of the Basin was covered by seasonally dry

forests at this time. This simulation highlights the

possibility that the C3 forest signal between 17 and 914C ka BP in southern Amazonia (De Freitas et al.,

2001) may signify deciduous rather than evergreen

forest. Further support comes from LGM model

simulations by Cowling et al. (2001) which also show

an increase in tropical seasonal forest at the expense of

tropical rainforest in central and south-eastern Ama-

Table 2

Land surface carbon storage (vegetation biomass and soil organic

matter) in the Amazon basin as a proportion of total carbon storage

by the terrestrial biosphere

Time interval Carbon

storage

in the

Amazonian

basin

Carbon

storage in

the entire

terrestrial

biosphere

% contribution

of the Amazon

basin

Pre-industrial

(1900–1910

average)

273.6 2466 11.1

Mid-Holocene (1) 222.5 1363 16.3

Mid-Holocene (2) 227.1 1466 15.5

Mean 224.8 1414.5 15.8

Last glacial

maximum (1)

140.1 562 24.9

Last glacial

maximum (2)

128.8 931 13.8

Mean 134.5 746.5 18.0

All figures are in Gt C. Amazonian numbers from Table 1, global

numbers for the palaeo-simulations from Beerling (1999).

zonia. Cowling et al.’s simulations further suggest that

changes in vegetation structure may have been at least

as important as changes in biome type at the LGM.

Their model results simulate a 34% decrease in Basin-

average leaf-area index (i.e. canopy density) relative to

modern values, due to decreased photosynthetic rate,

forced by a decrease in atmospheric CO2 from 360 to

200 Amol mol�1. Irrespective of the composition or

structure of forest communities that occupied the

Amazon Basin at this time, it is nevertheless clear

that both the model and the palaeo-data refute the

dglacial refugia hypothesisT advocated by Haffer

(1969), Prance (1982), and Whitmore and Prance

(1987), which states that most of the Amazon Basin

was covered by savanna at the LGM.

Net primary productivity (NPP) estimates are

difficult to compare directly with evidence from the

fossil record, but it is clear that the cool, low CO2

environment of the LGM severely restricted vegeta-

tion carbon uptake and lowered NPP relative to the

Holocene (Fig. 2d–f). We note that the increase in

productivity of Amazonian vegetation between the

Mid-Holocene and Pre-Industrial era (Fig. 2d–e),

coincident with the CO2 rise from 280 to 300 ppm,

is consistent with the observed rise in above ground

biomass shown by tropical forests in long-term

monitoring studies (Phillips et al., 1998). Basin-wide

changes in NPP (Fig. 2d–f) closely track changes in C

storage in vegetation biomass and soil organic matter

(Fig. 3), particularly in regions of evergreen and

deciduous broad-leaved forests, which have the

highest potential for carbon accumulation in trunk


biomass by virtue of their longevity (Chambers et al.,

1998).

There is a general trend in the model simulations of

increasing cover of evergreen broad-leaf forests (rain

forests), total C storage and NPP in the Amazon Basin

between the LGM, the Mid-Holocene, and the Pre-

Industrial period (Fig. 1 and Fig. 2, Table 1 and Table

2), which reflects the effects of climatic amelioration

and rising atmospheric CO2 levels on NPP. There is a

significant increase in evergreen broad-leaf forests at

the expense of deciduous forests (Fig. 2b,c), and a

67% increase in total Amazon C storage from the

LGM (135 Gt C) to the Mid-Holocene (225 Gt C)

(Fig. 3, Table 2). There is a further marked increase

in simulated evergreen forests and a 22% increase in

total C storage between the Mid-Holocene (225 Gt C)

and the Pre-Industrial era (274 Gt C) (Table 2). These

Holocene changes (albeit for only two time slices) are

in line with the palaeo-data discussed above, includ-

ing the precipitation records from the Bolivian

Andes.

It is interesting to note that although the size of the

Amazon C sink at the LGM is simulated to be 50%

smaller than in Pre-Industrial times, the relative size of

this C store at the LGM, by comparison with the total

C storage in the entire terrestrial biosphere, is

simulated to be nearly double that for the Pre-

Industrial era (Table 2). It appears, therefore, that

Amazonia played an even greater role in the global C

budget during the last ice age than today. This can

most likely be attributed to the fact that much of the

high latitude temperate deciduous forests and boreal

forests were largely replaced by tundra and continen-

tal ice-sheets. It is also noteworthy that the importance

of C storage in soil organic matter relative to the total

land surface C budget of Amazonia was markedly

higher at the LGM (70%) than in the Mid-Holocene

(50%) and Pre-Industrial period (40%) (Fig. 3d–f,

Table 1). The latter can be attributed to the cool LGM

climate (ca. 5 8C lower than present) (Stute et al.,

1995) slowing decomposition rates and allowing litter

accumulation.

It should be emphasized that none of these

simulations take account of the feedback between

vegetation and climate, a feature which, although

poorly understood, is of particular importance for the

Amazonian hydrological cycle. For example, a

simulated reduction in Amazon forest cover at the

LGM has been found to significantly reduce evapo-

transpiration and consequently tropical precipitation

(Levis et al., 1999). Clearly, understanding and

quantifying such feedbacks remains an important goal

for vegetation and climate modellers.

4. Comparison between Amazon carbon storage

fluxes and atmospheric carbon dioxide and

methane fluxes

4.1. Comparison with the ice-core CO2 record of

Taylor Dome, Antarctica

Antarctic ice-core records (e.g. Monnin et al.,

2001) show that atmospheric CO2 concentrations rose

from ca. 190 to 270 ppmv (parts per million by

volume) during the transition from the LGM to the

beginning of the Holocene (ca. 11 cal ka BP). A

continuous Holocene CO2 record from Taylor Dome

(Antarctica) (Indermuhle et al., 1999) showed that

CO2 concentrations decreased from 268 ppmv at 10.5

cal ka BP to 260 ppmv at 8.2 cal ka BP, and that over

the subsequent 7000 years CO2 concentrations

increased almost linearly to ca. 285 ppmv by 1 cal

ka BP.

The sharp rise in atmospheric CO2 levels during

the last glacial-Holocene transition must have been

driven by changes in ocean C storage, rather than

fluxes in the terrestrial C store, since this rise in CO2

coincided with an increase in global terrestrial C

storage (vegetation and soils) of 668 Gt C (estimated

by the SDGVM; Beerling, 1999), due to expansion of

terrestrial ecosystems, which would have actively

sequestered atmospheric CO2. Therefore, the absolute

amount of C transferred to the atmosphere during the

last deglaciation must have been even higher than that

revealed by the ice-core records.

Indermuhle et al. (1999) used inverse carbon cycle

modelling to suggest that, contrary to the last glacial-

Holocene transition, the observed Holocene CO2

variations were driven in large part by changes in

global terrestrial biomass. They postulated that the

terrestrial biosphere sequestered 110F47 Gt C

between 11 and 7 cal ka BP and released 195F40

Gt C between 7 and 1 cal ka BP. Of this 195 Gt C, ca.

145 Gt C were modelled to have been taken up by the

oceans and 50 Gt C taken up by the atmosphere,


equivalent to the observed 25 ppm rise in atm CO2

concentration over this time.

Adams and Faure (1998) reconstructed past

changes in the global geographic extent of the

dominant terrestrial ecosystems from previously

published palaeoecological data and estimated total

terrestrial ecosystem C storage based on summation of

estimates of living (biomass), dead (litter) and soil C

components determined from closest modern ana-

logues. Using this approach, they estimated that the

terrestrial biosphere sequestered ca. 195 Gt C between

9 and 6 cal ka BP, which is in line with the estimate by

Indermuhle et al., 1999 of 110 Gt C based on inverse

carbon cycle modelling and the Taylor Dome ice-core

data.

The Amazonian palaeoecological data discussed

above suggest that Amazonia played no role in this C

sequestration over most of this time period, but

instead acted as a significant net C source, since the

available data indicate that following the Late-glacial

period there was savanna expansion at the expense of

forests and increased fires in the early-mid Holocene.

Northern hemisphere ecosystems must therefore have

sequestered most of this C, since there is well-

documented palaeoecological evidence for pole-ward

expansion of boreal forests and expansion of tropical

ecosystems into the Sahel and Sahara of northern

Africa during this time (e.g. Wright et al., 1993;

Adams and Faure, 1998).

Numerous independent estimates have been made

of regional and global terrestrial carbon storage fluxes

between ca. 7 and 1 cal ka BP, using a variety of

approaches. Adams and Faure (1998) estimate from

palaeodata that there was a net loss of only 27 Gt C

from the total terrestrial biosphere between 6 and 0 cal

ka BP. Carbon losses of 3, 19.1 and 27 Gt C have been

estimated using a similar approach for Europe (Peng et

al., 1994), China (Peng and Apps, 1997), and Siberia

(Monserud et al., 1995) respectively. A combination of

palaeoecological data synthesis and biome/biosphere

modelling by Beerling (2000) and Indermuhle et al.,

1999 have produced estimates of 24–30 Gt C loss for

northern Africa. Clearly, none of these estimates

accounts for the magnitude of terrestrial C loss (195

Gt C) that has been hypothesized. If this hypothesized

C loss is correct, other terrestrial regions of the world

must have supplied the missing C. However, this C

loss coincided with the dominant phase of global

peatland expansion between 5 and 2 ka BP (reviewed

by Gorham, 1991), which would have constituted a

major C sink, since modern peatland ecosystems of the

taiga are estimated by Gorham (1991) and Adams and

Faure (1998) to store 465 Gt C. This peatland C sink

must have been a significant offset to any terrestrial C

sources.

Given that present-day Amazonian tropical forests

are estimated to account for about 10% of the carbon

stored in terrestrial ecosystems (Melillo et al., 1993),

any changes in Amazonian biomass between 7 and 1

cal ka BP must have had an important effect upon the

net terrestrial C budget. As discussed above, the

SDGVM predicts an increase in Amazon C storage

(vegetation biomass plus soil organic matter) of

between 46.5 and 51.1 Gt C between the mid-

Holocene and the present day (Table 2) (corroborated

by the palaeoecological data), which, together with

the increase in high-latitude peatland C storage over

this time, would have more than compensated for the

terrestrial C losses, resulting in a net terrestrial C sink

rather than C source between 7 and 1 ka BP. This

implies that, contrary to Indermuhle et al.’s hypoth-

esis, the oceans must have been the key contributor

to the steady atmospheric CO2 rise between 7 and 1

ka BP, and not the terrestrial biosphere. This

inference is supported by Broecker et al. (1999,

2001), although Brovkin et al. (in press) conclude

that more proxies are needed to determine the relative

contributions of terrestrial C vs. oceanic C to this

Holocene CO2 rise.

4.2. Comparison with polar ice-core methane records

Atmospheric methane is another important compo-

nent of the global carbon cycle and has been shown

by ice-core studies to exhibit even greater Late

Quaternary fluxes than atmospheric CO2. During

deglaciation methane concentrations doubled from

LGM values of 350 ppbv to Late-glacial values of ca.

700 ppbv (Chappellaz et al., 1990, 1993). The Late-

glacial period was interrupted by a sharp decrease in

concentrations to ca. 500 ppbv during the Younger

Dryas chronozone (11–1014C ka BP, ca. 13–11.5 cal

ka BP) (Brook et al., 1996), while the Holocene

interglacial was also characterized by methane varia-

tions of 15%. Blunier et al. (1995) analysed the GRIP

ice-core record and showed high methane concen-


trations of ca. 725 ppbv from 11.5 to ca. 8 cal ka BP,

followed by low concentrations (600–625 ppbv) until

ca. 3 cal ka BP, after which there was a steady rise to

pre-industrial levels of 725 ppbv.

Although it is generally agreed that, prior to

anthropogenic influences, the dominant methane

source was natural wetlands, the relative contributions

from the mid-high northern hemisphere latitudes

versus the tropics have been open to question and

the subject of considerable debate (e.g. Chappellaz et

al., 1990; Street-Perrott, 1992). However, Chappellaz

et al. (1993) point out, by reference to their high

resolution methane record, that the large increase in

methane concentrations during the Late-glacial inter-

stadial (ca. 14.5–12.7 cal ka BP), took place when

much of the northern hemisphere wetlands were still

covered by ice sheets, lending support to an Ama-

zonian origin for much of this methane. Furthermore,

Maslin and Burns (2000) reconstructed the outflow

history of the Amazon river over the last 14,000 years

and showed there was a 60% decrease in Amazon

discharge during the Younger Dryas chronozone (ca.

12.7–11.5 cal ka BP), signifying widespread decrease

in Amazonian wetlands, correlating with a sharp

decrease in atmospheric methane concentrations. This

further supports the hypothesis that the Late-glacial

methane fluxes were driven in large part by changes in

Amazonian wetlands.

It is also noteworthy that the GRIP Holocene

methane curve correlates extremely well with the

Amazon precipitation record inferred from the suite of

palaeo-data discussed above. Pollen, stable carbon

isotope, and charcoal records show maximum Hol-

ocene aridity (i.e. lowest wetland cover) in lowland

Amazonia between ca. 9 and 3 cal ka BP (Fig. 1) and

increasing precipitation over the past 3 millennia.

These trends also match the Lake Titicaca water-level

records. It is therefore tempting to infer from these

close correlations that fluxes in geographic extent of

Amazonian wetlands were a key determinant in

atmospheric methane variability, not just through the

Late-glacial period, but the Holocene as well.

However, Ruddiman and Thomson (2001) dis-

agree. They argue convincingly from several lines of

evidence that the Late Holocene methane increase

cannot be explained by any dnaturalT ecosystem flux

(either from tropical wetlands or high latitude peat-

lands), but instead by an anthropogenic source from

rice cultivation over the last 5000 years. Their

strongest argument for an anthropogenic origin comes

from their observation that this late Holocene pre-

Industrial methane increase is a feature unique to the

present interglacial and that no similar pattern is

evident over the last 400,000 years (Chappellaz et al.,

1990; Petit et al., 1999).

5. Conclusions

A review of recently published palaeoecological

data allows reconstruction of changes in Amazonian

ecosystems, and hence carbon storage since the LGM.

Contrary to Haffer’s (1969) hypothesis, it is now clear

that Amazonia was predominantly forested at the

LGM, although there is evidence for savanna expan-

sion at the margins, and modern biogeographic data

from disjunct species distributions raises the possi-

bility that much of the Basin at this time may have

been covered by seasonally dry forests rather than rain

forests. These interpretations are supported by model

simulations (SDGVM forced with two GCM climate

models) which confirm that most of Amazonia was

forested at the LGM and that southern Amazonia was

covered by deciduous broad-leaf forests. Although we

estimate Amazon C storage at the LGM to be only

135 Gt C (50% smaller than Pre-Industrial estimate),

most likely due to C limitation caused by low

atmospheric CO2 levels, its proportion of C storage

in the entire terrestrial biosphere was found to be

almost twice that for the Pre-Industrial era.

The model shows that between the LGM and the

Mid-Holocene there is a significant increase in ever-

green broad-leaf forests at the expense of deciduous

forests and a 67% increase in total Amazon C storage,

attributable to rising temperatures and atmospheric

CO2 levels.

Although palaeoecological data show that the

Amazon Basin was dominated by rain forests

throughout the Holocene, there is clear evidence

from ecotonal areas near the northern and southern

margins of the Basin for increased extent of savannas

and fires, and therefore reduced precipitation, in the

Early-mid Holocene compared with the Late Hol-

ocene. This is also in line with our model simulations

which show a further increase in rain forests and a

22% increase in total C storage between the Mid-


Holocene (225 Gt C) and the Pre-Industrial era (274

Gt C).

Consideration of these changes in the Amazon C

sink through the Holocene in the context of palaeo-

data from other parts of the world suggests that the

terrestrial biosphere as a whole acted as a net C sink

during this time, and that contrary to Indermuhle et al.

(1999) hypothesis, the observed CO2 rise from 260

ppmv at 8.2 cal ka BP to 285 ppmv at 1 cal ka BP

(revealed by the Antarctic Taylor Dome ice-core

record) was driven by release of C from the oceans

rather than land.

Close similarity between the pattern of Amazon

ecosystem changes and the Greenland methane ice-

core record implies that Amazonia played an impor-

tant role in driving atmospheric methane fluxes over

the last glacial–Holocene transition, through changes

in extent of its wetlands and rain forests.

Acknowledgements

We thank the following: Hugues Faure, Bruno

Turcq, and Luiz C.R. Pessenda for the invitation to

present this paper at the session on dCarbon Cycle

ChangesT (sponsored by IGCP 404 and INQUA) at

the 31st International Geological Congress, Rio de

Janeiro, Brazil, 2000; the Royal Society and the

University of Leicester for providing financial support

for FEM to attend this conference; R. Pollington for

cartographic assistance with Fig. 1; funding for DJB

provided by a Royal Society University Research

Fellowship and the Leverhulme Trust; Herman

Behling and Mark Bush for useful comments on an

earlier draft of the paper.

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