( Springer-Verlag 1999
Climate Dynamics (1999) 15 : 29—42
T. Goslar · B. Wohlfarth · S. Björck · G. Possnert
J. Björck
14
Variations of atmospheric
Dryas transition
C concentrations over the Allero/ d-Younger
Received: 15 April 1998/Accepted: 9 July 1998
Abstract Highly variable atmospheric radiocarbon
concentrations are a distinct feature during the last
deglaciation. The synchronisation of two high-resolution AMS 14C-dated records, Lake Gościa9 z5 , and
a floating Late Weichselian glacial varve chronology at
the Aller+d-Younger Dryas transition allowed us to
assess in detail atmospheric *14C changes between late
Aller+d and early Preboreal. The combined data set
shows a drastic rise in *14C during the first 200 years
or so of Younger Dryas and the two following about
500 year- long 14C plateaux. Model experiments which
included variations in the geomagnetic field, atmospheric CO variations and a drastic reduction in
2
North Atlantic Deep Water flux at the onset of
Younger Dryas allowed to reproduce the distinct rise in
*14C during the first 200 years of Younger Dryas fairly
well. Also the drop in *14C at the Younger
Dryas/Holocene boundary seems reasonably explained
by changes in North Atlantic Deep Water circulation.
However, the reason behind the anomalous behaviour
of the *14C signal in the middle of Younger Dryas
remains an open question.
T. Goslar ( )
Radiocarbon Laboratory, Institute of Physics,
Silesian Technical University, Krzywoustego 2 PL- 44 100 Gliwice,
Poland
E-mail: [email protected]
B. Wohlfarth
Department of Quaternary Geology, Lund University,
Tornavägen 13, S-223 63 Lund, Sweden
S. Björck
Geological Institute, University of Copenhagen,
"ster Voldgade 10, DK-1350 Copenhagen K, Denmark
G. Possnert
Tandem Laboratory, Uppsala University, Box 533,
S-75121 Uppsala, Sweden
J. Björck
Department of Quaternary Research, Stockholm University,
S-106 91 Stockholm
1 Introduction
The transition period between the Last Glacial Maximum and the present Interglacial was characterised by
highly variable atmospheric radiocarbon (14C) concentrations. These are expressed by several long-lasting
radiocarbon plateaux, sudden drops in radiocarbon
ages (e.g. Hughen et al. 1998; Björck et al. 1996; Goslar
et al. 1995a) and by a large offset between radiocarbon
and calendar years (Kitagawa and van der Plicht 1998;
Bard et al. 1993). A radiocarbon calibration curve
based on U/Th-dated corals has been proposed to
extend the Holocene dendro-calibration curve back in
time (Stuiver and Reimer 1993). Unfortunately, the
U/Th data points published so far (Bard et al. 1993,
1996; Edwards et al. 1993) still seem too scarce to reveal
atmospheric *14C variations in great detail. These
problems might, however, be overcome with AMS
14C dating of laminated sediment records. The new
data sets from the Cariaco basin and Lake Suigetsu
provide now, e.g. possibilities to assess radiocarbon
variations more in detail (Hughen et al. 1998) and as far
back as 35 14C ky BP or 38.0 cal ky BP (Kitagawa and
van der Plicht 1998).
Lacustrine terrestrial sequences have the advantage
that radiocarbon and varve ages can directly be related
to interstadial and stadial vegetation changes. One of
these records is the laminated sediment sequence from
Lake Gościaz5 in Poland, which covers a time period
from late Aller+d up to the middle Holocene (Goslar
et al. 1995a).
The most unequivocal way of matching different
radiocarbon-dated terrestrial chronologies to each
other is to use the beginning and/or end of a radiocarbon plateau as a time-synchronous marker horizon.
Since variations in past atmospheric 14C/12C ratios can
be regarded as time-synchronous world wide, the rapid
changes in radiocarbon ages at the end/beginning of
a radiocarbon plateau are ideal markers for correlating
30
between densely 14C-dated records. Here, we combine
three different high-resolution (AMS 14C) dated terrestrial data sets, which all show a distinct drop in
radiocarbon ages at the Aller+d (AL) — Younger Dryas
(YD) boundary: the laminated lake record from Lake
Gościa9 z5 (LG) in Poland (Goslar et al. 1995a), the nonlaminated, but high-resolution AMS 14C-dated lake
sediment sequence from lake Madtjärn (LM) in SW
Sweden (Björck et al. 1996) and, a floating late Weichselian glacial varve chronology, which forms part of the
Swedish Time Scale (STS) (Fig. 1). The synchronisation
between two of these different types of high-resolution
dated terrestrial chronologies (LG and the STS) at the
AL—YD boundary, allows the possibility of assigning
calendar-year ages to the floating varve chronology
and makes it feasible to reconstruct and discuss variations in atmospheric *14C (*14C!5.) over this
transition.
Wiggles in the radiocarbon calibration curve are
equivalent to variations of past atmospheric 14C concentrations. Rapid drops in 14C age and plateaux of
constant 14C age reflect, respectively, increases and
decreases of atmospheric 14C. Following models by e.g.
Hughen et al. (1998), Stocker and Wright (1996), Björck
et al. (1996) and Goslar et al. (1995a), the maximum in
atmospheric 14C concentration around the Aller+d-
Fig. 1A Location of Lake
Gościa9 z5 (¸G) in Poland, Lake
Madtjärn (¸M) and the Swedish
Time Scale. B The different
regional varve chronologies
which form part of the STS are
shown by striped bars; those of
Cato (1987), Lidén and Cato (in
preparation) and Strömberg
(1985, 1989) cover the Holocene
part, while those of Brunnberg
(1995), Strömberg (1994),
Kristiansson (1986) and
Ringberg (1991) make up the
Late Weichselian part. Mullsjön
(MS) and the 770-y long varvedclay chronology (open rectangle)
discussed in the text are
indicated
Goslar et al.: Variations of atmospheric 14C concentrations
Younger Dryas transition could be interpreted as
a manifestation of a drastic weakening of the North
Atlantic thermohaline circulation (THC). Such
a change in THC in turn may be a plausible mechanism
of rapid global climatic changes (Broecker 1991). Goslar et al.’s (1995) and Björck et al.’s (1996) conclusions
were based on a few 14C dates only. By combining the
three records from LG, LM and the STS and by synchronising the data sets from LG and the STS, we are
here able to document the large maximum in *14C
during the early YD by a more extensive set of dates.
2 The data set
2.1 Lake Gościa9 z5 (LG)
Lake Gościa9 z5 is situated in central Poland (Fig. 1A) and has a sequence of annually laminated sediments deposited between about
3 and 12.8 cal ky BP (Goslar et al. 1995a, b; Ralska-Jasiewiczowa et
al. 1992). The large set of proxy climate data allows a detailed
correlation to the major European pollen zones and to the Greenland ice core record (GRIP). By ‘wiggle-matching’ the densely
spaced AMS 14C measurements on terrestrial macrofossils to the
Holocene calibration curve, it was possible to tie the floating varve
chronology from LG to calendar years (Goslar et al. 1995b).
Combined pollenstratigraphic investigations and stable 18O
analyses allowed defining the transitions between AL and YD to
Goslar et al.: Variations of atmospheric 14C concentrations
31
12580$130 cal y BP, and between YD and Preboreal (PB) to
11440$120 cal y BP (Goslar et al. 1995a). Recently, Goslar and
Ma9 dry (1998) presented an improvement of the wiggle-matching
technique, and updated the calendar-year age of the LG chronology,
through the match to the revised oak chronology (Björck et al.
1996). Accordingly, the ages of the AL/YD and YD/PB transitions have been adjusted to 12 650$60 cal y BP and
11 510$40 cal y BP.
2.2 Lake Madtjärn (LM)
Madtjärn is a small lake situated in southwest Sweden (Fig. 1A).
Pollen stratigraphy, macrofossil and stable isotope analyses allow
an assignment to the Scandinavian Late Weichselian and early
Holocene climatic and vegetational development (Björck and Digerfeldt 1991; Wohlfarth et al. in preparation). LM became deglaciated
during AL and the marine isolation occurred slightly before the
AL-YD boundary. Late AL sediments comprise a succession from
marine clays to gyttja clay, clayey gyttja and moss-rich clayey gyttja.
The transition from clay gyttja to clayey fine detritus gyttja marks
the boundary between YD and PB as a distinct lithological change.
The densely spaced AMS 14C dates performed on terrestrial macrofossils give a high-resolution 14 °C stratigraphy for the AL-YD-PB
pollen zones (Björck et al. 1996, 1997).
2.3 The Swedish Time Scale (STS)
Glacial and postglacial varved clays are a common feature in lakes
and peat bogs along the Swedish east coast, in the estuaries of the
large rivers in northern Sweden and in the Baltic Sea. The glacial
varved clays deposited during the Late Weichselian and Early Holocene retreat of the Scandinavian inland ice and reflect the seasonal
melting of the ice sheet with their silty summer and clayey winter
laminae (Björck et al. 1992). The postglacial varved clays cover the
Holocene time period up to present and are mainly delta sediments
deposited in the estuaries of the large rivers in northern Sweden
(Widerlund and Roos 1994; Cato 1987). Already early this century
varved clays were used to create a time scale for the ice recession from
southern to northern Sweden, by correlating successively younger
varve diagrams with each other (De Geer 1912). This time scale has
later been revised and many more varve diagrams were added to
strengthen the Holocene (Lidén and Cato in preparation; Cato 1987;
Strömberg 1985) and the Late Weichselian part of the chronology
(Brunnberg 1995; Strömberg 1994, 1989; Ringberg 1991; Kristiansson 1986) (Fig. 2A, B). Based on these revisions and on the correlations presented by Lidén and Cato (in preparation), Brunnberg
(1995), Strömberg (1994, 1989, 1985), Björck and Möller (1987), Cato
(1987) and Kristiansson (1986), the STS had been regarded as a continuous calendar year chronology covering the last about 13300
varve years (Wohlfarth et al. 1995). Accordingly, earlier AMS
14C dates obtained on terrestrial macrofossils from the Late Weichselian varves were related to varve years BP by assuming that the STS
is a correct, calendar-year time scale (Wohlfarth et al. 1995, 1993).
Recently, however, several pieces of evidence have been put forward, which clearly indicate errors in the STS.
1. The calibrated age of an AMS 14C date obtained on terrestrial
macrofossils from the Holocene varves differs by several hundred
years compared to the corresponding varve age and indicates an
error between approximately 2000—5000 varve y BP (Wohlfarth
et al. 1997).
2. A discrepancy of several hundred years exists between the varve
age BP and the calibrated age BP for one of the most widely
spread marker horizons in varved and non-varved sequences
along the Swedish east coast, the Yoldia ingression (Björck et al.
1996). This ingression which is synchronous with the Preboreal
Oscillation (11300—11150 cal y BP) (Björck et al. 1997) is dated to
Fig. 2 A The linkage of the different regional varve chronologies of
the Swedish Time Scale according to Björck et al. (1992), Strömberg
(1994), Brunnberg (1995) and Björck and Möller (1987). The connection between Strömberg’s (1989) and Lidén and Cato’s (in preparation) chronologies is based on drainage varves. Brunnberg’s
chronology (1995) is in turn linked to Strömberg’s (1989) chronology
by overlapping varve diagrams and the marker varves of the Yoldia
ingression (marked by a striped bar). The match between Brunnberg’s
(1995) and Kristiansson’s chronology (1986) is based on a colour
change visible in both chronologies and on overlapping varve diagrams. The connection between Kristiansson’s (1986) and Ringberg’s
(Ringberg 1991) chronology is a tentative link, suggested by Björck
and Möller (1987). B Location of errors within the regional varve
chronologies: '300 y in the middle Holocene and possibly about
300 y in the early Holocene (Wohlfarth et al. 1997; Björck et al. 1996),
#47 y within Kristiansson’s (1986) chronology and about # 1400 y
in the early Late Weichselian (Wohlfarth in preparation). The addition of 47 y could be confirmed through the new 770-year long varve
chronology, while the prolongation by 1400 y is still tentative and has
to be evaluated by further investigations. Pollenstratigraphic investigations on the varves in south easternmost Sweden (Ringberg 1991)
by Ising (in preparation) show that these varves were deposited during
the B+lling and Older Dryas pollen zone. In comparison with the
GRIP ice core (Johnsen et al. 1992), Ringberg’s (1991) chronology is,
therefore, placed between about 14000 and 14600 calendar y BP
about 10430—10310 varve y BP (Brunnberg 1995) (Fig. 2A), which
indicates an offset of 870 y.
3. The AL-YD transition, which had been estimated in the varved
clays to about 12 varve ky BP is 600 years younger, when
Goslar et al.: Variations of atmospheric 14C concentrations
32
compared to the LG record (about 12.65 cal ky BP) (Goslar et al.
1995a; Wohlfarth et al. 1995).
4. An additional offset is visible during the older part of the STS, i. e.
the B+lling to early AL varves, when the varve ages BP of the
corresponding AMS 14C dates are compared to the 14C-U/Th
coral record (Wohlfarth 1996).
These pieces of evidence show beyond doubt, that the present-day
STS can not be regarded as a continuous calendar-year chronology
and that the critical parts have to be revised (Fig. 2B) before the Late
Weichselian AMS 14C dates can unequivocally be tied to varve
years BP.
All AMS 14C dates obtained so far can be divided into three parts
(Table 1, Fig. 1B).
A. The youngest 14C dates were obtained from the site Mullsjön,
which is situated in south-central Sweden. The varve diagram
could be correlated to the oldest part of Strömberg’s (1994) varve
chronology. The varve ages assigned here to the 14C dates were
set 800 y older than Strömberg’s (1994) varve years, according to
the discrepancy outlined already.
B. The local varve ages attributed to the AMS 14C dates from
the sites Nedre Emmaren, Tynn, Gummetorpasjön, Hargsjön,
Adlerskogssjön and Glottern are based on an independent
770-year long varve chronology, which has been established
in southeastern Sweden by cross correlating 26 varve diagrams
(Wohlfarth et al. in print) (Figs. 1B, 2B). In order to avoid
the introduction of a new set of local varve years, this
new chronology was connected to Kristiansson’s (1986)
varve chronology at the local varve year 1940. Following
this match the new chronology covers the local varve years
1701—2475.
C. The sequences from Toregöl and Lillsjön are beyond the range of
the 770-year long chronology. Therefore, no local varve ages
could be attributed to the radiocarbon dates obtained from these
two sites.
Table 1 AMS 14C dates from the glacial varves and their estimated calendar year age BP. The calendar-year ages assigned to group A were
obtained by adding 800 years to Strömberg’s (1994) chronology and those in group B by synchronising the records with the data set from
Lake Gościa9 z5 (see text for further explanation). The calendar-year ages assigned to group C are only tentative. L"leaves, Lf"leaf fragment,
F"flower, S"seeds, UP"unidentified plant remains, *"plant remains possibly reworked, d"uncertain age caused by fungi (see
Wohlfarth et al. 1995). Samples marked with * and d and are not included in Figs. 3 and 4
Lab. No.
Locality
Local
varve years
Macrofossils
submitted for AMS measurement
AMS 14C
yr BP
Estimated
cal yr BP
A
Ua-4212
Ua-2741
Ua-4214
Mullsjön
Mullsjön
Mullsjön
11010$10
11027$54
11115$11
10,160$115
9,640$190
10,170$195
11,810
11,827
11,915
Ua-4215
Mullsjön
11178$50
10,140$155
11,978
Ua-2742
Ua-4216
Ua-4217
Mullsjön
Mullsjön
Mullsjön
11356$25
11437$19
11471$14
Salix herbacea (Lf ), Salix indet. (Lf )
?Dryas oct. (Lf )
Salix indet.. (Lf; B), Salix/Betula (Lf ),
Betula nana (Lf )
Salix indt. (Lf ) , Salix herbacea (Lf ), Betula nana (Lf )
Salix/Betula (Lf ), Oxyria (S), Caryophyllaceae (S)
Betula/Salix (Lf ), brown mosses (L, F, stems)
Salix indet. (Lf ), Betula nana (Lf ), Oxyria (S)
Salix herbacea (Lf )
9,945$115
10,620$155
10,330$175
12,156
12,237
12,271
B
Ua-4496
Ua-10180
Glottern
Glottern
1856$50
1872$66
10,585$465
11,550$300*
12,506
12,522
Ua-10187
Ua-2544
Ua-10186
Ua-4359
Ua-11527
Ua-4493
Ua-10185
Gummetorpasjön
Hargsjön
Gummetorpasjön
Hargsjön
Hargsjön
Adlerskogssjön
Gummetorpasjön
1938$4
1938$2
1968$25
1973$31
1990$24
2003$59
2009$16
10,420$220
11,405$145*
11,040$110
10,610$110
10,384$130d
10,830$165
11,230$100
12,588
12,588
12,618
12,623
12,640
12,650
12,659
Ua-2753
Ua-10184
Ua-4358
Ua-10183
Ua-11518
Ua-3131
Ua-10182
Ua-11234
Hargsjön
Gummetorpasjon.
Hargsjön
Gummetorpasjon
Tynn
Tynn
Gummetorpasjön
Nedre Emmaren
2010$45
2044$16
2055$50
2090$18
2125$35
2125$35
2138$30
2146$23
10,480$150
10,970$90
10,980$100
11,030$120
10,511$130d
10,890$120
11,470$130
10,885$250
12,660
12,694
12,705
12,740
12,775
12,775
12,788
12,796
Ua-10181
Ua-11233
Gummetorpasjön
Nedre Emmaren
2199$32
2221$52
11,450$240
10,745$240
12,849
12,871
C
Ua-4635
Ua-4634
Ua-2750
Ua-2752
Ua-4945
Lillsjön
Lillsjön
Toregöl
Toregöl
Lillsjön
11,895$860*
12,320$470*
11,520$225
11,820$150
11,530$130
?
?
13,557?
13,557?
13,563?
?
?
?
?
?
Salix/Betula (Lf )
Empetrum (L), Salix indet (Lf ), Salix/Betula (Lf ),
Dryas oct. (L), Arenaria (S), insects
Dryas oct. (L), Salix/Betula (Lf )
Dryas oct. (Lf, stem), ¸eguminosa (S)
Dryas oct. (L), Salix/Betula (Lf )
Dryas oct. (¸), Salix polaris (L)
Salix indet. (Lf )
Salix polaris (L), Dryas oct. (L, F)
Salix reticulata (L), Salix polaris (L),
Salix indet, (Lf ), Ericaceae (L), Dryas oct. (L)
Betula/Salix (Lf ), UP
Salix/Betula (Lf ), Dryas oct. (L)
Salix polaris (L)
Dryas oct. (L), Betula/Salix (Lf )
Salix/Betula (Lf ), Dryas oct. (Lf, F)
Salix indet. (Lf ), Dryas oct. (L, F), insects, UP
Ericacea (L, F), Salix/Betula (Lf ), Dryas oct. (L)
Salix indet. (Lf ), Dryas oct. (L, S), Arenaria (L),
Salix polaris (L)
Saxifraga (F), Salix indet. (Lf ), Silene (L)
Salix indet. (Lf ), Salis polaris (L), Dryas oct. (L, F),
Betula nana (L)
Salix polaris (L), Salix/Betula (L), Dryas oct. (L)
Salix polaris (L), Armeria, Dryas oct. (L)
Dryas oct. (Lf ), Betula/Salix (Lf ), UP
Dryas oct. (Lf ), Betula/Salix (Lf ), UP
Salix/Betula (Lf )
Goslar et al.: Variations of atmospheric 14C concentrations
33
The only possibility to relate the floating AMS 14C-dated varve
chronologies of the STS to an absolute time scale, is by synchronising the data set with other, well-dated absolute records
at certain marker horizons, such as the end or beginning of
a 14C plateau.
3 Synchronisation of the data set
In the pollen record from LG the AL-YD transition
is marked by a decrease in tree pollen, an increase
in Artemisia and Juniperus, and by a drastic decline
in the algae ¹etraedron minimum, which in turn
gives indications for declining summer temperatures
(Goslar et al. 1993; Ralska-Jasiewiczowa et al. 1992).
A similar interpretation can be derived from the decrease in the d18O record of authigenic carbonates,
which coincides with the vegetational changes at
12650$60 cal y BP (Goslar and Ma9 dry 1998; Goslar
et al. 1995a). At LM the boundary between AL and YD
is marked by a decrease in tree pollen (Betula alba,
Pinus), an increase in shrubs (Betula nana, Juniperus),
herbs (Artemisia) and decreasing pollen concentrations
(Björck et al. 1996).
A significant feature in both, the LG and the LM
record is the rapid drop in radiocarbon ages at the
transition from AL to YD, which is followed by a long
radiocarbon plateau (Fig. 3 and inset figure). The same
marked drop in 14C ages from 11—10.8 14C ky BP to
10.6—10.4 14C ky BP is also visible in the AMS 14C dated
770-y long sequence of the Swedish varves at around
the local varve year 2000. Based on the assumption that
the atmospheric 14C/12C content was the same worldwide, we can use this rapid decrease in radiocarbon
ages as a basis for correlating the floating varve sequence to LG (Fig. 3). Indirect support for this correlation is given by pollen stratigraphic investigations
performed on the 770-y long local varve chronology
which show a distinct YD pollen signal starting between the local varve years 1850—2000 (J. Björck submitted). The AMS 14C dates obtained on the varved
sequence from Lake Mullsjön, show indications for
a transition between the first and second YD plateau
(Björck et al. 1996; Goslar et al. 1995a), i.e. the drop in
14C ages from about 10.6—10.4 14C ky BP to about
10.1—10 14C ky BP. By fitting the 14C ages from
Mullsjön to the LG curve a difference of about 800 y is
apparent when compared to the varve ages BP in
Strömberg’s (1994) chronology. Based on this fit and
based on the apparent offset of 870 y (Björck et al. 1996,
1997), we tentatively added a minimum of 800 varve
years to Strömberg’s (1994) chronology in order to
account for the difference in the younger part of
the STS. The calendar-year assignment of the 14C ages
obtained from Lillsjön and Toregöl, i.e. from the
non-revised part of Kristiansson’s (1986) chronology
remains, however, tentative until the new 770-y long
chronology has been extended.
Fig. 3 Radiocarbon ages versus calendar-year ages during the Late
Weichselian and early Holocene. The records of Lake Gościa9 z5 (solid
circles) (Goslar et al. 1995a) and the Swedish varved clays (open
diamonds) were synchronised by the method described in the text.
Open triangles represent the 14C-U/Th dates on corals from Barbados and Mururoa Atoll (the calculated mean) (Bard et al. 1993),
Huon Peninsula (Edwards et al. 1993) and Tahiti (Bard et al. 1996).
Crosses denote 14C dates derived from varved sediments from the
Cariaco Basin (Hughen et al. 1998). All data sets are presented with
a double standard deviation. The smooth line represents a spline
function fitted to all the data, except to those from the Cariaco
Basin. The inset figure shows the radiocarbon-depth record from
Lake Madtjärn (open circles) (Björck et al. 1996) and the rapid drop
in 14C ages at the Aller+d/Younger Dryas boundary
The synchronised records, LG and the STS, (Fig. 3)
are rather consistent with the 14C-U/Th coral data
(Bard et al. 1993, 1996; Edwards et al. 1993). A generally good agreement exists also between the combined
set of the Swedish, Polish and coral data, the recently
published 14C dates from varved marine sediments
from the Cariaco Basin (Hughen et al. 1998) as well as
the record from Lake Suigetsu (Kitagawa and van der
Plicht 1998). All records indicate a marked drop in
14C age from 11—10.9 to 10.6—10.4 14C ky BP, completed within 100—200 y and followed by two distinct,
about 500 y long plateaux at approximately 10.4 and
10 14C ky BP. The indication of a long radiocarbon
plateau at 10 14C ky BP, which is followed by a shorter
plateau at 9.5 14C ky BP and by a gradual decrease of
radiocarbon ages after 10.8 cal ky BP, is in excellent
agreement with the (not shown in Fig. 3) results of the
high-precision dated German pine chronology (Björck
et al. 1996; Kromer and Becker 1993). Although the
dates are rather sparse and reveal a fairly large scatter
between 13 and 12.7 cal ky BP, there are indications in
our data set for a short 14C plateau 200 y before the end
of AL. A similar radiocarbon plateau of 200—300 y is
34
also visible in the Cariaco basin and Lake Suigetsu
data sets (Hughen et al. 1998; Kitagawa and van der
Plicht 1998).
The drop in 14C ages at 12.65 cal ky BP, which
coincides in LG (Goslar and Ma9 dry 1998; Goslar et al.
1995a) and LM (Björck et al. 1996) with the transition
from AL to YD occurs in Lake Suigetsu at approximately the same time, i.e. around 12.7 cal ky BP
(Kitagawa and van der Plicht). An age of 12.65 cal ky
BP for this transition agrees also well with the estimate
of about 12.7 ky BP indicated by the GRIP ice core
(Johnsen et al. 1992). The recently published records
from the Cariaco basin (Hughen et al. 1998) suggest,
however, a distinctly older placement for the AL/YD
transition and for the drop in radiocarbon ages
(12.9—13.0 ky BP). In this data set, the duration of the
YD period of 1330 y is around 200 y longer than the
1150 y consistently indicated by European sites (Goslar
et al. 1995a, b; Hajdas et al. 1995, 1993) and the GRIP
ice core (Björck et al. 1996; Johnsen et al. 1992). On the
other hand, the timing of the YD in the Cariaco basin
agrees with that recorded in the GISP2 ice core
(Hughen et al. 1998). It remains to be solved whether
the apparent non-synchronism of the AL/YD
transition in different sites is an artefact of inadequate
calendar-year chronologies or whether it reflects the
real dynamic behaviour of the global climate system. At
any rate, a delay between climatic changes at the two
nearby sites on the Greenland summit (GISP and
GRIP) seems fairly unlikely. Independently of the
meaning of the Cariaco AL-YD boundary, our records
indicate that the AL-YD transition occurred in Europe
at around 12.65 cal ky BP, just at the marked drop in
14C age.
4 Possible causes of atmospheric D 14C variations
Natural radiocarbon is one of the most robust tracers
of oceanic and atmospheric circulation, because it is
produced by cosmic rays in the atmosphere only (Lal
1992) and then distributed to other reservoirs involved
in the global carbon cycle, i.e. the oceans and the
biosphere. The total amount of 14C on Earth is thus
controlled by the balance between 14C production and
its radioactive decay in all reservoirs. The concentration in particular reservoirs depends, on the other
hand, on reservoir size and rate of carbon exchange
between them. 14C concentration in the atmosphere is
higher than in any other reservoir, since the source of
radiocarbon is in the atmosphere itself. It is obvious
that a weakening of the carbon exchange between other
reservoirs will produce an increase in *14C!5.. As demonstrated by Stuiver et al. (1991) and Stuiver and Braziunas (1993), the major features of the *14C!5. curve in
the Holocene may be explained by variations in the
14C-production rate alone, controlled by the strength
Goslar et al.: Variations of atmospheric 14C concentrations
of the geomagnetic field and by solar activity. In fact,
there are no clear indications for significant changes in
the global carbon cycle during the Holocene. Another
situation concerns the Late Weichselian. During this
period, changes in carbon circulation are manifested by
a distinct rise of atmospheric CO content, as evid2
enced in studies of Antarctic and Greenland ice cores
(Anklin et al. 1997; Jouzel et al. 1992; Leuenberger et al.
1992; Barnola et al. 1991, 1987; Neftel et al. 1988). Also,
firm evidence exists for a drastic change in the intensity
of the vertical mixing of oceanic water between the Late
Glacial Maximum (LGM) and the Holocene.
The ventilation of the modern deep oceans is driven
mostly through sinking of large surface water masses in
the North Atlantic, where North Atlantic Deep Water
(NADW) is formed and in the belt around Antarctica,
where formation of Antarctic Bottom Water (AABW)
occurs. As shown by many paleoceanographic studies
(e.g. Sarnthein et al. 1995, 1994; Keigwin and Lehman
1994; Charles and Fairbanks 1992; Keigwin et al. 1991;
Jansen and Veum 1990; Boyle 1988; Duplessy et al.
1988; Boyle and Keigwin 1987), the formation of
NADW was much weaker during the LGM than today. Changes in the carbon cycle, implied by Antarctic
and Greenland ice cores and marine studies, should
have a distinct imprint on the pattern of *14C!5. variations during the Late Weichselian. Indeed, the amplitude of the *14C!5. change observed during early YD
(Fig. 4) is much larger (50—70 per mille) and longer
(500—700 y) than any *14C!5. fluctuation during the
following 12500 calendar years (see inset in Fig. 4). This
*14C!5. maximum could be plausibly explained by
a cessation of NADW flux. According to general circulation models (GCM), shutting down NADW formation and reducing northward heat transport from
equatorial regions would have a profound effect on
North Atlantic regional and perhaps global climate
(Manabe and Stouffer 1997, 1995, 1988; Rahmstorf
1995). While a much weakened NADW flux during the
LGM is indeed evidenced by paleoceanographic studies, the question of reduced NADW formation during
the YD is, on the other hand, still under discussion (e.g.
Keigwin and Lehman 1994; Keigwin et al. 1991;
Charles and Fairbanks 1992; Jansen and Veum 1990).
However, there is growing evidence that the sinking of
water masses in the North Atlantic proceeded also
during YD, but to shallower depths than today (Sarnthein et al. 1994; Lehman and Keigwin 1992).
The possible connection between reduced NADW
formation and an increase of *14C!5. during early YD
has already been discussed (Hughen et al. 1998; Stocker
and Wright 1996; Björck et al. 1996; Goslar et al.
1995a). The residual *14C, obtained after removal of
the long-term trend from the observed record, was
compared in these works with a *14C!5. simulated
using several simple models of the carbon cycle (Bard
et al. 1994; Broecker et al. 1990; Keir 1988; Oeschger
et al. 1975) or in a model including a more realistic,
Goslar et al.: Variations of atmospheric 14C concentrations
Fig. 4 Atmospheric radiocarbon concentrations versus calendar-year
ages during the Late Weichselian and early Holocene. The concentrations are expressed as per mil deviation (*14C) from the standard. The
symbols correspond to those in Fig. 3. Some of the outlying data
from those shown in Fig. 3 have been omitted. The maximum of
*14C during the early Younger Dryas is striking. The smoothed line
represents a spline function fitted to all data, except to some outlying
data. For comparison, some of the largest fluctuations of *14C
during the Holocene are shown in the inset on the same scale. These
fluctuations correspond to the periods: (a) 9.6—9.25 cal ky BP
(Kromer and Becker 1993); (b) 7.3—7.05 cal ky BP (Stuiver and
Reimer 1993) and (c) 2.8—2.2 cal ky BP (Pearson et al. 1993)
zonally averaged ocean circulation (Stocker and Wright
1996). In most of those simulations the *14C wiggles
were calculated for several scenarios of oceanic ventilation changes while other parameters of the carbon cycle
and the 14C-production rate were kept constant.
Here, we present another approach, which simultaneously takes into account a few factors that could
influence the varying atmospheric radiocarbon concentration during YD: (1) long-term variations of
the 14C production rate, driven by the well-reconstructed changes of the geomagnetic field; (2) terrestrial and
oceanic mechanisms of the observed changes of atmospheric CO content and, (3) possible fluctuations of
2
the NADW flux.
5 Model simulations of observed D 14C
5.1 Model description
As a basis for the simulations we used the PANDORA
model (Broecker et al. 1990) of the global carbon cycle.
Oceanic circulation of carbon in this model is represented through the fluxes of water, organic matter and
35
Fig. 5 A Schematic representation of oceanic water circulation in
the PANDORA model of the global carbon cycle (after Broecker et
al. 1990). In the steady-state of pre-industrial ocean the fluxes of
individual loops are as follows: L1, 6 Sv; L2, 13.5 Sv; L3, 10.5 Sv; L4,
4 Sv; L5. 3 Sv; L6, 15 Sv; L7, 20 Sv; L8, 100 Sv. The additional loop
L9 (0 Sv in the pre-industrial steady-state) has been introduced to
represent the upper part of the NAIW flux. B Variations of fluxes of
individual loops of oceanic water circulation, assumed in the scenarios S2, S3 and S4 of the model simulations discussed in the text.
The variations of loops L1 and L2 were the same in all scenarios. The
scenarios for loops L3 and L9 are distinguished by line thickness (S2:
thin line, S3: medium line, S4: thick line)
carbonates between particular well-mixed boxes. The
circulation of water (Fig. 5A) has the form of a few
loops (L1—8), which connect the surface boxes with the
deep oceans. All paraeters of oceanic circulation have
been adopted from Broecker et al. (1990). The atmosphere (600 Gt C) exchanges carbon with the oceans
(total 37000 Gt C in the pre-industrial steady state, 1 Gt
C"1012 kg of carbon) and the terrestrial biosphere
(2150 Gt C). The carbon fluxes between the surface
oceanic boxes and the atmosphere are proportional to
the partial pressures of CO in the surface water and in
2
the air. The standard flux (for pCO "280 latm) is
2
equal to 15 mol/m2y. The atmosphere-biosphere fluxes
are proportional to the mass of carbon stored in appropriate reservoirs. The formation of NADW is represented in the model by the sum of the fluxes L1, L2 and
L3 (30 Sv altogether, 1 Sv"106 m3/s). This is distinctly higher than the NADW flux of about 20 Sv in the
real ocean (Schmitz and McCartney 1993). The model
fluxes have been set (Broecker et al. 1990) to reproduce
36
correctly the distribution of 14C in the pre-industrial
ocean. In fact, when no transport of organic matter and
carbonates exists, 14C is transported to the deep ocean
due to both the large scale advective circulation (such
as NADW) and the turbulent diffusion. Therefore, the
exchange of carbon between reservoirs is always stronger than that attributed to water currents only.
The PANDORA model includes parameters, which
control and potentially could be responsible for Late
Weichselian changes of atmospheric CO content.
2
These are (1) the alkalinity of the surface water, (2) the
strength of the biological pump, which depends on the
residence time of phosphorus in the surface reservoirs
and, (3) the mass of the terrestrial biosphere.
Calculations of *14C!5. during the period 15—10
14C cal ky BP (Fig. 6) were made in a few runs: (1) with
all model parameters kept constant, except for the
14C production rate dependent on the geomagnetic
field; (2) with the 14C production as in (1), but combined with the parameters controlling atmospheric
CO to reproduce the observed variations of CO ; (3)
2
2
with all parameters as in (2) and a variable flux of
NADW.
Goslar et al.: Variations of atmospheric 14C concentrations
5.2 14C production rate
Variations of the 14C-production rate were derived
from paleomagnetic data of Tric et al. (1992) and McElhinny and Senanayake (1982), using the relations of
Lal (1988). As pointed out by Goslar et al. (1995a),
calculated 14C concentrations depend on the initial
value of *14C and on the model adjustment of the
absolute value of 14C production. In the standard
model, the 14C production, which corresponds to the
current geomagnetic field, maintains the steady state
*14C!5."0&. However, using such a production and
initiating the run from the actual *14C level at 15 cal
ky BP, we obtain too high *14C!5. values over the
whole Holocene. The best fit of the simulated to
the observed Holocene *14C!5. curve is obtained when
the 14C production is systematically lowered by 2%
(Goslar 1996). With such an adjustment, variations of
the geomagnetic field almost completely explain the
general decline of *14C!5. between 15 and 10 cal ky BP
(Fig. 6), but they are too weak to produce the large
*14C maximum during the early YD.
5.3 Size of the atmospheric carbon reservoir and carbon
exchange between atmosphere and ocean
Fig. 6 Results of modelling the atmospheric 14C concentrations
during the late Weichselian and early Holocene. The crosses represent the same data set as in Fig. 4. The results of the model simulations are shown by the following lines: dashed line, *14C variations
driven only by fluctuations of the geomagnetic field; dotted line,
*14C variations driven by fluctuations of the geomagnetic field and
including the mechanisms responsible for changes of atmospheric
CO ; dashed-dotted line, the scenario as mentioned plus a reduced
2
NADW flux to 7 Sv during the whole YD (scenario S1); solid line,
the same scenario but with the NADW flux changing gradually at
the boundaries of YD (thin line: scenario S2, medium line: scenario
S3, thick line: scenario S4). The details of all scenarios are given in the
text
Another mechanism responsible for *14C!5. variations
are changes in atmospheric CO . Studies performed on
2
Antarctic and Greenland ice cores indicate a rather
linear, slow increase of CO from about 200 ppm at 17
2
ky BP to about 280 ppm in the Holocene. This change
could produce the decrease of *14C!5. by 25—35& (Lal
and Revelle 1984; Keir 1983, Siegenthaler et al. 1980).
Goslar et al. (1995a) argued that the gradual increase of
CO could not be responsible for the rapid variations
2
of *14C!5. during YD. Nevertheless, the incorporation
of the CO scenario allowed a better reproduction of
2
the general slope of the *14C curve between 15 and 10
cal ky BP (Fig. 6).
The rate of CO exchange between the mixed layer of
2
the oceans and the atmosphere is another parameter
affecting *14C!5.. A decrease of the ocean-atmosphere
CO flux by 50% rises *14C!5. by about 20& (e.g.
2
Siegenthaler and Beer 1988). Björck et al. (1996) suggested, that during YD this flux might have been 50%
weaker than today, possibly caused by abundant cover
of sea ice and icebergs as shown from paleoceanographic studies (Haflidason et al. 1995; Koi Karpuz
and Jansen 1992). The recent simulations of Stocker
and Wright (1996) show, however, that freezing the
Atlantic and Pacific oceans north of 45°N would raise
the *14C!5.content only by 2&. A distinct increase of
*14C!5. (by 20&) would only occur after freezing the
Southern Ocean (south of 47.5 °S). Although the question whether this ocean warmed or cooled during YD is
partly still under debate (Mabin 1996; Lowell et al.
1995; Manabe and Stouffer 1995; Denton and Hendy
Goslar et al.: Variations of atmospheric 14C concentrations
1994), there is growing evidence for an anti-phase relationship between the Southern and Northern Hemispheres during YD (Blunier et al. 1997). Moreover,
steeper thermal gradients during the cold YD period
may have resulted in stronger trade winds (Hughen et
al. 1998; Bard et al. 1994; Mayewski et al. 1993), an
effect which could enhance CO exchange even by
2
a few tens of per cent. Having in mind that those
opposite mechanisms partly cancel each other, we regard the effect of altered air-sea exchange during the
YD as still unknown.
5.4 Including vertical ocean circulation
The most probable mechanism of the rapid *14C!5.
increase during early YD is the drastic reduction of
vertical oceanic circulation. In previous papers several
scenarios of circulation were considered.
1. The scenario where NADW flux was reduced from
30 Sv to 10 Sv during the whole YD (Goslar et al.
1995a) gave the highest *14C!5. values at the end of
YD, in contradiction to experimental data. A similar
peak of *14C!5., accompanying the drastically reduced North Atlantic overturning through the
whole length of YD, was also reproduced by Stocker
and Wright (1996).
2. Björck et al. (1996) considered a reduced eddy diffusivity of the deep oceanic reservoir by 50% during the
first 200 y of YD as an effect of a sudden fresh-water
forcing and a gradual recovery to modern values
until the onset of the Holocene. Such a scenario
reproduces relatively well the observed shape of the
YD *14C!5. maximum. GCM models (Rahmstorf
1995; Mikolajewicz and Meier-Reimer 1994), however, predict a gradual decrease of the NADW flux
as a response to gradually increased freshwater forcing. The successive decrease of freshwater forcing
would then lead to an abrupt switch-on of the
NADW circulation, which contradicts with Björck
et al.’s (1996) scenario.
3. Another approach (Hughen et al. 1998) assumed
that NADW formation ceased completely during
YD and was replaced by a gradually increasing flux
of North Atlantic Intermediate Water (NAIW).
NAIW could serve as an effective sink for 14C!5., but
would have heated the North Atlantic region much
less than NADW (Rahmstorf 1994).
In the simplest scenario presented here (hereinafter
denoted as S1), NADW flux instantly reduced at the
onset (12650 cal y BP) and restored at the termination
(11510 cal y BP) of YD. The 14C concentrations obtained in the later part of that period are too high and
the rise of *14C!5. at the beginning of YD seems too
slow and slightly delayed with respect to the observed
development (Fig. 6). Furthermore, the simulated decline of *14C!5. at the beginning of the Holocene distinctly lags the experimental one.
37
Recent GCM calculations by Manabe and Stouffer
(1997) suggest, that NADW flux and sea surface temperatures could change with a phase lag. Manabe and
Stouffer (1997, hereinafter denoted MS97) supplied an
extra 0.1 Sv of freshwater to the North Atlantic surface
during a period of 500 y. They obtained an initial
(200 y) rapid decrease of the thermohaline circulation
(THC), which was followed by a distinctly slower decrease during the following 300 y. When the freshwater
forcing had ceased, the THC re-intensified gradually
within about 250 y and reached temporarily even higher values than before the disturbance. The inactive
THC appeared thus rather unstable and could restore
itself, in contrast to earlier suggestions (Rahmstorf
1995). However, the results obtained by MS97 seem
more reliable, since their atmospheric circulation
model is more realistic than the highly parametrised
model applied by Rahmstorf (1995). Nevertheless the
duration of the increase in THC in the MS97 model is
a few times shorter than that assumed by Björck
et al. (1996).
In our scenario (S2) we therefore, used Manabe and
Stouffer’s (1997) results for the simulations of *14C!5..
The significant drop in sea surface temperature in the
MS97 experiment started about 100 y after the beginning of the decline in THC and the main cooling was
completed during the following 100—150 y. The duration of this cooling is similar to that recorded in the
sediments of Lake Gościa9 z5 , i.e. 150 y between 12725
and 12575 cal y BP. We, therefore, prescribed the beginning of the MS97 THC decline to the year 12820 cal
y BP and adopted the declining part of the MS97 curve
(years 0—500) for our scenario of NADW flux changes
around the AL/YD boundary. Similarly, we dated the
middle part of the main MS97 warming (year 700 in the
MS97 time scale) to 11510 cal y BP and used the rising
part of the MS97 curve (years 500—1250) for the
NADW scenario around the YD/PB boundary. Unfortunately, the period of freshwater supply in the MS97
simulations was significantly shorter than the real
duration of YD and the THC curve in the MS97
experiment could not cover the middle part of YD. We,
therefore, assumed that NADW flux was constant during this intermediate period (12320—11710 cal y BP).
The resulting scenario of the variations in deep water
flux (S2), which were used as an input in the PANDORA model, is shown in Fig. 5B (loops L1—2 and L3).
As discussed before, the deep ocean water fluxes in the
steady-state pre-industrial model are about 50% higher
than the large-scale fluxes in the real ocean. Accordingly, the amplitude of the NADW fluxes in our simulations is proportionally higher than that of the THC
circulation in MS97.
The increase of *14C!5. obtained by the MS97-like
scenario S2 occurs earlier and fits the reconstructed
maximum in *14C!5. at the YD onset better than the
simplified scenario S1 (Fig. 6). In addition, the earlier
decline of the calculated *14C!5. values around the
38
YD/PB boundary appears closer to the observed one.
This seems to support the indication of MS97, that
during the YD cold spell, NADW flux did not change
instantly and not exactly in phase with climatic changes. However, *14C!5. values in the S2 simulation
remained rather constant after 12.3 cal ky BP, in contradiction to the distinct drop of *14C!5. documented
by our reconstructions.
The decline of *14C!5. in the mid- and late YD could
be reproduced well by Hughen et al. (1998), who allowed for changes in NAIW flux in their simulations.
The comprehensive study of d13C distributions in the
North Atlantic (Sarnthein et al. 1994) suggests ventilation of the deep Atlantic almost as strong as now
during the YD, although North Atlantic surface waters
penetrated to shallower depths than today. According
to Rahmstorf (1994), NAIW heats the North Atlantic
region much less effective than NADW circulation.
However, a simple replacement of NADW by NAIW
flux produces a very slow increase of *14C!5., which is
much weaker than that observed at the AL/YD boundary. Also Manabe and Stouffer (1997) did not obtain
a significant increase of NAIW in their studies. One
possibility to reconcile these contradictory data is, to
assume that NAIW flux was weak during a short period at the beginning of Younger Dryas. Such a short
(350 y) break in ventilation might not have had an
appreciable influence on deep oceanic d13C. If the
water circulation were to cease completely, the present
flux of organic matter from the North Atlantic surface
(about 1.4 Gt C/y in the PANDORA model) might
decrease the deep Atlantic d13C with a rate of 0.3&/
century. However, the biological productivity at the
surface rapidly decreases as soon as the circulation has
ceased, because it is limited by the availability of nutrients, which are supplied from the deep reservoirs. Indeed, the organic carbon flux integrated through the
years 12650—12300 in our S2 run is only 127 Gt C. This
amount could not alter the d13C of Atlantic water by
more than 0.3&. In fact, the d13C values of deep Atlantic water during the time slice 10.35—10.8 14C ky BP are
0.2—0.3& lower than between 8.35—9.1 ky 14C BP (see
e.g. Figs. 16 and 18 in Sarnthein et al. 1994). A strengthening of NAIW a few hundred years after the beginning
of YD does neither contradict the results obtained by
Manabe and Stouffer (1997), who did not prolong their
simulations of inactive THC state beyond the year 500.
The scenario of deep Atlantic circulation, particularly weakened during early YD, seems concordant
with high-resolution palaeoclimate records from the
Greenland ice cores (Johnsen et al. 1992; Alley et al.
1993) and from Lake Gościa9 z5 (Goslar et al. 1995a). It is
also in agreement with a number of palaeoclimatic
reconstructions of the YD period in Europe, where the
coldest climate occurred during the early part of YD
(e.g. Isarin 1997; Renssen 1997; Walker 1995; Lowe et
al. 1995; Wohlfarth et al. 1994). Furthermore, the data
set from core ENAM 93—21 in the Faroe-Shetland
Goslar et al.: Variations of atmospheric 14C concentrations
Channel (Rasmussen et al. 1996) suggests that a strong
freshwater forcing in the North Atlantic occurred only
during the earliest 250—400 y after the onset of YD
(Björck et al. 1996).
Hughen et al. (1998) assumed a complete cessation of
both NADW and NAIW at the beginning of YD. We
simulated the appearance of a strong NAIW flux after
12300 cal y BP, while at 12200 cal y BP, the sum of
NADW and NAIW fluxes is assumed to be equal to the
pre-YD and Holocene NADW flux. Around the
YD/PB boundary, the NAIW flux gradually decreased,
just to keep the sum of NADW and NAIW fluxes
constant. Such a scenario is different from that proposed by Hughen et al. (1998), who assumed a slow
increase of NAIW through '1000 y. In general, GCM
simulations have, up to now, given either almost instantaneous increases of NADW-NAIW fluxes or
a gradual increase (e.g. MS97). However, this gradual
increase never lasted longer than 250 y. None of the
GCM simulations has yet dealt with a 1000 year long
period of increasing NADW-NAIW, which renders
Hughen et al.’s (1998) scenario unlikely.
The NAIW flux was represented first by the additional loop (L9 in Fig. 5A, B), which connects the
low- and high-latitude surface boxes of the Atlantic
(scenario S3). As the mean depth of the Atlantic surface
reservoirs is 1000 m, loop L9 seems too shallow to
represent NAIW well. In effect, the drop of *14C!5.,
after the activation of L9 (Fig. 6), is quickly restrained
when all surface Atlantic reservoirs become saturated
with atmospheric radiocarbon.
In an improved scenario (S4), NAIW was represented by a combination of loops L3 and L9. L3 connects
the surface Atlantic reservoirs with the deep Atlantic
but, similar to the pattern of shallower NADW flow
obtained by Rahmstorf (1994, Fig. 1b), it does not
reach the Arctic and Indo-Pacific oceans. We assumed
that the L3 flux increased between 12300—12200 cal
y BP to the same value as before YD (Fig. 5B). The L9
flux changed accordingly, in order to keep the sum of
NADW and NAIW fluxes equal to the pre-YD
strength of NADW. This scenario implies that half of
the water flowing out from the North Atlantic reservoir
penetrated deeper than 1000 m, which is again in reasonable agreement with the pattern obtained by Rahmstorf (1994). Such a scenario reproduces the observed
*14C!5. quite well. Nevertheless, it is not easy to distinguish whether the scenario with a gradual NAIW
increase (Hughen et al. 1998), or the scenario with
a rapid intensification of NAIW (scenario S4, this
work), allow reproducing better the decline of *14C!5.
during the YD.
The decline of *14C!5. around the end of YD can
now be explained almost perfectly. This is mainly due
to the simulation by Manabe and Stouffer (1997), who
showed that the increase of THC flux could be initiated
well before the major warming. Although scenario S4
reproduced the *14C!5. drop at the YD/PB transition
Goslar et al.: Variations of atmospheric 14C concentrations
fairly well, it should be treated with caution, since it is
not supported by GCM data. However, if started from
a level of about 200 & (which was really the case), this
drop is well explained by Manabe and Stouffer’s (1997)
THC model, independently of the mechanisms, which
lead to the starting level.
Some simulations of YD-like changes of NADW
circulation were also published earlier by Sakai and
Peltier (1996) and Stocker and Wright (1996). In these
studies, NADW flux was stopped during the period of
meltwater pulses MWP-IA and MWP-IB (Bard et al.
1996; Fairbanks 1989). The *14C!5. maxima produced
by these scenarios, however, offset the *14C!5. maximum by about 1000 y.
The appearance of a stronger penetration of the deep
Atlantic by Antarctic Bottom Water and a gradual
weakening of the meridional overturning in the IndoPacific ocean were two additional features of the MS97
simulation. The connection between the deep Antarctic
and the Atlantic is in our model represented by the loop
L4. Doubling the strength of the loop leads to an
increase of *14C!5. of less than 0.5&. The Indo-Pacific
overturning (I-PO) in the MS97 simulations gradually
decreased, in phase with the decrease of the Atlantic
THC flux (between the years 0—500). In fact, the decrease of I-PO during that period was incorporated in
our scenarios S2—4 and was represented by the
weakened loop L1. However, this decrease continued
until the year 800—900, i.e. when Atlantic circulation
had already been restored. The decrease of I-PO
around the YD/PB boundary produces again a significant (about 200 y) lag between the observed and the
simulated drop of atmospheric radiocarbon (not shown
in Fig. 6). One can, however, imagine that the minimum of I-PO could still occur during a period of weak
Atlantic THC, if the freshwater forcing in the MS97
simulation were kept sufficiently long.
6 Discussion and conclusions
The synchronisation of a floating AMS 14C-dated glacial varve chronology with the high-resolution AMS
14C-dated terrestrial data sets from Lake Gościa9 z5 at
the AL-YD transition (12.65 cal ky BP), lead to a dense
record of *14C!5. changes covering the time period
between late AL to early PB. The combined data set
shows in general good agreement to other high-resolution chronologies, such as the 14C-U/Th coral series
(Bard et al. 1996, 1993; Edwards et al. 1993), the
Cariaco and Lake Suigetsu varve record (Hughen et al.
1998; Kitagawa and van der Plicht 1998) and the German pine chronology (Kromer and Becker 1993). A distinct feature evident in our combined records is the
drastic rise in *14C!5. during the first 200 y of YD,
followed by two about 500 y long 14C plateaux. A less
distinct rise in *14C!5. is evidenced some 200 y before
the end of AL, which leads into a shorter, around 200 y
39
long 14C plateau. However, since the data set is still too
sparse for the time period before 12.7 cal ky BP, we
focus here on modelling the marked *14C!5. rise during early YD.
While variations in the geomagnetic field explain
almost entirely the general decline of *14C!5. between
15 and 10 cal ky BP, they are too weak to reproduce
the observed *14C!5. maximum in our model. Including atmospheric CO variations in the model run, the
2
general slope of the *14C!5. curve between 15 and
10 cal ky BP can be simulated better, although the
*14C!5. maximum at the onset of YD is not reproduced. The most likely mechanism for drastic changes
in *14C!5. is a significant reduction in vertical ocean
ventilation (Stocker and Wright 1996; Björck et al.
1996; Goslar et al. 1995a). We, therefore, performed
model simulations including the two parameters mentioned and assumed gradual changes of NADW flux
near the boundaries of YD, following those obtained in
recent GCM simulations by Manabe and Stouffer
(1997).
Following Manabe and Stouffer’s (1997) and Stocker
and Wright’s (1996) model experiments for the YD, the
formation of deep water in the North Atlantic is terminated or greatly reduced, when the input of melt
water from the decaying ice sheets exceeds 0.1 or 0.2 Sv.
This in turn results in a decreased uptake of 14C by the
ocean and an increase in *14C!5.. As shown by Bard et
al. (1996), the large pulse of melt water from the
decaying ice sheets (Fairbanks 1989) occurred as much
as 1000 y prior to the YD cooling. The implied long
delay in response of the North Atlantic heat conveyor
to the freshwater forcing could not yet be successfully
modelled. Also the maximum of *14C!5. accompanying the weakening of NADW expected during melt
water pulse MWP-IA (Sakai and Peltier 1996; Stocker
and Wright 1996), would precede the real maximum by
about 1000 y.
Alternatively, Björck et al. (1996) proposed that the
rapid slow-down of the thermohaline circulation at the
AL-YD transition could have been triggered by a sudden drainage of large dammed glacial lakes on both
sides of the North Atlantic. This is concordant with
reconstructions by Clark et al. (1996), implying that
the main source of water for MWP-IA was in Antarctica, which, according to Manabe and Stouffer’s
(1997) model, could not effectively disturb the NADW
circulation.
The magnitude and extent of the early YD *14C!5.
maximum clearly exceeds the smaller *14C maximum
in our records (about 200 y before the end of AL) and
the wiggles seen in LG during YD and PB. The same
picture emerges from the record of ice rafted debris
(IRD) (Rasmussen et al. 1996; Björck et al. 1996).
Clearly minor IRD peaks are associated with smaller
*14C!5. maxima, while the large IRD peak during early
YD seems synchronous with the large maximum in
*14C values. This implies that smaller *14C!5. peaks
40
might be triggered by increased meltwater input
into the North Atlantic due to a rapid decay of the ice
sheets during warm periods, which lead to a temporary
shut-down of deep-water convection and to a cooling
(Björck et al. 1996).
We conclude that the large *14C!5. maximum during early YD could be reproduced in our model by
combining variations in the geomagnetic field, in atmospheric CO and a significant short-term reduction in
2
vertical ocean ventilation. The more than 1000-y long
*14C!5. decline during YD and PB, however, could be
reproduced only roughly. The almost unrealistic
change in atmosphere-ocean flux postulated by Björck
et al. (1996) or rather ad-hoc proposed scenarios involving NAIW fluxes (Hughen et al. 1998, and this
work) to explain this decline, suggest that the underlying reason for this anomalous behaviour of the *14C!5.
signal is still enigmatic.
Acknowledgements BW would like to thank Tomas Nilsson, Markus Olsson, Björn Holmquist, Siv Olsson and Dan Hammarlund for
their help during field work, Felicia Dobos and Jessica Ademark for
sieving the macrofossils from the Swedish varved clays and Gina
Hannon for helping with plant macrofossil determinations. Comments by T. Stocker, E. Bard and an anonymous reviewer helped to
improve the manuscript. BW’s work was supported by the Swedish
Geological Survey (SGU) and the Swedish Natural Science Research
Council (NFR) through grants 03—500/92, 03—792/93:9, G-GU
10197—300 and G-AA/GU 10197—301. The work of TG was sponsored by the Polish Committee of Scientific Research through grant
6 P04E 027 10.
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