Introduction
The Western Siberian Lowland (WSL) can be considered as one of the most
vulnerable permafrost-bearing territories with respect to ongoing climate
change, due to (i) the dominance of discontinuous, sporadic and intermittent
permafrost coverage rather than continuous and discontinuous permafrost of
central and eastern Siberia and the Canadian High Arctic, (ii) its flat
area and high impact of flooding and thermokarst development, and, most
importantly, (iii) its high stock of ancient and recent organic carbon in
the form of partially frozen peat deposits. Due to the importance of the
boreal and subarctic continental zones in the Earth's carbon cycle and the
high vulnerability of circumpolar zones to climate warming, the majority
of conducted works have been devoted to the biogeochemistry of organic carbon
and sediments in large rivers of the Russian boreal circumpolar zone (Gordeev
et al., 1996, 2004; Moran and Woods, 1997; Lobbes et al., 2000; Dittmar and
Kattner, 2003; Gebhardt et al., 2004; Cooper et al., 2008; Nikanorov et al., 2010a, b;
Holmes et al., 2000, 2001, 2012; Pokrovsky et al., 2010; Feng et al., 2013).
While these studies have allowed for the quantification of the carbon and
major element delivery fluxes from the continent to the Arctic Ocean, the
mechanisms responsible for carbon and metals mobilization from the
soil/groundwater to the rivers remain very poorly understood. The WSL offers
a unique site to test various hypotheses of element sources and to reveal
related mechanisms as it presents the full gradient of the permafrost
coverage, climate and vegetation over homogeneous sedimentary basement rock,
essentially peat soil, flat orography and similar annual precipitation.
Taking advantage of these features, in their pioneering studies, Frey et
al. (2007a, b) and Frey and Smith (2005) provided a first-order assessment of
the relative contributions of shallow surface water and deep groundwater to
small western Siberian rivers. Their study was conducted during the summer
baseflow season, presenting the largest contrast between permafrost-free and
permafrost-affected rivers. This allowed them to conclude that climate
warming should shift the permafrost-affected part of the region from
surface feeding to groundwater feeding, while the permafrost-free zone may
remain unaffected.
However, unlike many regions of the world, the boreal and subarctic river
regions exhibit extreme seasonal variations in discharge and chemical
elements concentrations (see Voronkov et al., 1966; Gordeev and Sidorov,
1993; Gordeev et al., 1996; Gislason et al., 1996; Gaillardet et al., 2003;
Rember and Trefry, 2004; Zakharova et al., 2005, 2007; Bagard et al., 2011,
2013; Prokushkin et al., 2011; Guo et al., 2004b, 2007; Olefeldt and Roulet,
2012; Voss et al., 2015). The quantitative description of these systems,
therefore, requires an understanding of how weathering rates and riverine
fluxes of major and trace elements as well as their main carrier (organic
carbon) vary seasonally. High seasonality implies significant variations in
the source of the elements in river flow over the year, which is further
accentuated by high variability in the depth of the active layer and relevant
contributions of mineral soil weathering and the leaching of the soil organic
horizon. As such, the chemistry of fluxes on the seasonal scale depends on
the relative role of mineral dissolution vs. plant litter (organic soil)
leaching. Although several recent studies have used isotopic techniques in an
attempt to resolve the sources of elements in subarctic rivers (Engström
et al., 2010; Keller et al., 2010; Pokrovsky et al., 2013a; Mavromatis et
al., 2014), the relative contributions of mineral and plant litter/organic
soil components remain poorly constrained, particularly for boreal watersheds.
The purpose of the present work is to improve our understanding of western
Siberian river transport of organic and inorganic carbon and major elements
(Ca, Mg, K, Si) via studying numerous watersheds across the 1500 km
latitudinal profile during three main hydrological seasons: winter baseflow,
spring flood and summer–autumn period (Zakharova et al., 2014). For a working
hypothesis, we assume, following the previous works of Frey et al. (2007a,
b), that the permafrost controls riverine chemical composition via regulating the
degree of (i) groundwater feeding and (ii) leaching of elements from unfrozen
(active) soil layers. Because groundwaters in the permafrost zone are
discharged to the river via unfrozen taliks underneath the river bed
(Anisimova, 1981; Bagard et al., 2011, 2013), it can be suggested that the
impact of groundwaters via taliks will be mostly visible in large rivers, as is also known from the geocryological studies of the WSL (Fotiev, 1989,
1991). As a result, the contrast in groundwater-related element concentration
between rivers of different latitude is expected to be the largest during
winter baseflow. This is especially true in the WSL, exhibiting highly
homogeneous, extremely flat topography and similar lithological cover (peat,
sand and silt). Therefore, the first objective of this study was to test the
effect of river size (watershed area) on inorganic river water components
across the permafrost gradient. The second objective was to assess the effect
of the permafrost coverage on DOC, DIC and its isotopic composition in rivers
during different seasons. Specifically, during spring flood, when the majority
of the soil layer is frozen, only surface flux should be important and the
concentrations should reflect the degree of DOC and element leaching from the
plant litter. The largest contrast between rivers of different size is
therefore expected in August, whereas the spring flood should exhibit the
lowest differences in terms of DOC transport by rivers of different climate
and permafrost zones. Finally, the third objective of this study was
first-order assessment of the major river constituent concentration across
the 2000 km latitudinal profile. Here, we expect – in accordance with a general
knowledge of DOC and major cation concentration and export fluxes dependence
on temperature, vegetation and permafrost distribution (White and Blum, 1995;
Dessert et al., 2003; Gaillardet et al., 2003; Millot et al., 2003; Oliva et
al., 2003; Smedberg et al., 2006; Frey and McClelland, 2009; Prokushkin et
al., 2011; Beaulieu et al., 2012; Tank et al., 2012a, b; Olefeldt et al.,
2014) – a gradual or stepwise decrease in all river water constituents
northward, from permafrost-free to discontinuous and continuous permafrost
zone. Verifying the correctness of these research statements should allow for
the quantitative prediction of the degree of river water composition
modification in response to changing environmental conditions, notably the
increase in the thickness of active (unfrozen) layer, increasing the winter
discharge and augmenting plant biomass and productivity.
Study site and methods
Geographical setting
The Western Siberian Lowland (WSL) is the world second largest flooding territory,
after the Amazon's Varzea. The rivers (mainly the tributaries of the Ob, Pur,
and Taz) drain Pleistocene sands and clays, covered by thick (1 to 3 m) peat
and enclosing three main zones of the boreal biome – taiga, forest–tundra and
tundra. Approximate coverage of studied territory by sand, peat and clay
deposits in the first 3 m soil layer is shown in Fig. 1. Note that the peat
is always dominant on the watershed divides and bog zones, whereas the sand is
abundant along the river valleys. Quaternary clays, sands, and silts ranging
in thickness from several meters to 200–250 m have alluvial, lake-alluvial
and, rarely, aeolian origin south of 60∘ N and fluvio-glacial and
lake-glacial origin north of 60∘ N. The older (i.e., Paleogene and
Neogene) rocks are rarely exposed on the surface and are represented by
sands, alevrolites and clays, where carbonate material is present as
concretions of individual shells (Geological composition of the USSR,
1958). The climate is humid
semi-continental with a mean annual temperature (MAT) ranging from
-0.5 ∘C in the south (Tomsk region) to -9.5 ∘C in the
north (Yamburg). The annual precipitation increases from 550 mm at the
latitude of Tomsk to 650–700 mm at Noyabrsk and further decreases to
600 mm at the lower reaches of the Taz River. The annual river runoff
gradually increases northward, from 160–220 mm yr-1 in the
permafrost-free region to 280–320 mm yr-1 in the Pur and Taz river
basins located in the discontinuous to continuous permafrost zone (Nikitin
and Zemtsov, 1986). A detailed description of physico-geography, hydrology, lithology and
soil can be found in earlier works (Botch et al., 1995; Smith et
al., 2004; Frey and Smith, 2005, 2007; Frey et al., 2007a, b; Beilman et al.,
2009; Vorobyev et al., 2015) and in our recent limnological and pedological studies (Pokrovsky et al., 2013b; Shirokova et
al., 2013; Manasypov et al., 2014, 2015; Stepanova et al., 2015). A detailed
map of studied region together with main permafrost provenances and river
runoff in the WSL is given in Fig. 1, and the list of sampled rivers grouped
by watershed size and season is presented in Table 1. Permafrost zonation
in the WSL shown in this figure is based on extensive geocryological work in
this region (Baulin et al., 1967; Gruzdov and Trofimov, 1980; Baulin, 1985;
Liss et al., 2001). The hydrological parameters of the WSL rivers are described in Supplement S1.
List of sampled rivers, their watershed area and annual runoff. The
codes under the months identify the sampling sites listed in Table S1 in the Supplement. The
annual runoff was calculated following the approach of Frey et al. (2007b) as
explained in Supplement 1.
Number on the map
Season
N
E
River
Watersheds,
Annual runoff,
km2
mm yr-1
June
August
October
February
1
RJ-1
R-1
RF1
56∘31′48′′
84∘09′44′′
Ob
423 100
207
2
RJ-3
R-3
RF2
56∘46′19.5′′
83∘57′35.7′′
Prud
61.5
44.8
3
RJ-2
R-2
56∘43′15.0′′
83∘55′35.1′′
Chybyr′
8.14
44.8
4
RJ-4; R-10
R-4
RF3
57∘06′39.2′′
83∘54′41.1′′
Shegarka
12 000
58.3
5
RJ-5
R-5
RF4
57∘19′20.7′′
83∘55′53.8′′
Brovka
320
63.4
6
BL-3
56∘54′39.1′′
82∘33′33.3′′
Cherniy Klyuch
32
168
7
BL-2
57∘02′23.75′′
82∘04′02.44′′
Bakchar
3197
96.1
8
RJ-6
RF5
57∘36′43.3′′
83∘37′02.1′′
Malyi Tatosh
302
63.4
9
RJ-7
R-7
RF6
57∘37′17.3′′
83∘31′53.3′′
Bolshoy Tatosh
1020
74.6
10
RJ-8
R-8
RF7
57∘52′26.8′′
83∘11′29.9′′
Chemondaevka
177
63.4
11
RJ-9
R-9
RF8
57∘58′45.7′′
82∘58′32.2′′
Sugotka
275
63.4
12
RJ-10
RA-23
RF9
58∘04′20.8′′
82∘49′19.7′′
Chaya
27200
95.6
13
RJ-11
RF10
58∘23′16.8′′
82∘11′39.0′′
Tatarkin Istok
58.6
33.4
14
RJ-12
R-12
58∘24′38.0′′
82∘08′46.0′′
Istok
12.3
127
15
RJ-13
RA-22
R-13
RF11
58∘26′06.9′′
82∘05′43.6′′
Shudelka
3460
211
16
RJ-14
R-14
RF12
58∘33′03.1′′
81∘48′44.3′′
Chigas
689
180
17
BL-9
58∘32′05.8′′
80∘51′26.8′′
Karza
473
148
18
BL-6
58∘37′29.9′′
81∘06′09.0′′
Sochiga
510
148
19
RJ-15
RA-21; BL-4
R-17
RF64
58∘42′34.5′′
81∘22′22.0′′
Parabel
25 500
131
20
RJ-58
BL-5
R-15
RF65
58∘40′46.5′′
84∘27′56.6′′
Vyalovka
117
127
21
RF63
58∘59′37′′
80∘34′00′′
Vasyugan
63 780
177
22
RF62
59∘41′01.6′′
77∘44′33.9′′
Kornilovskaya
190
133
23
RF61
59∘44′09.2′′
77∘26′06′′
Levyi Il'yas
253
133
24
RF60
60∘08′43′′
77∘16′53′′
Koltogorka
220
155.4
25
RF58
60∘30′19′′
76∘58′57′′
Sosninskii Yegan
732
199
26
RJ-16
BL-36
60∘40′28.8′′
77∘31′29.4′′
Ob
773 200
216
27
BL-35
60∘44′10.9′′
77∘22′55.9′′
Medvedka
7
173
28
BL-34
60∘45′58.5′′
77∘26′12.6′′
Saim
26
173
29
BL-33
60∘47′29.3′′
77∘19′13.5′′
Mishkin Saim
32
173
30
BL-32
60∘49′32.3′′
77∘13 46.3′′
Alenkin Egan
44
173
31
BL-31
60∘50′43.6′′
77∘05′03.0′′
Kaima
31
173
32
BL-30
60∘55′41.0′′
76∘53′49.3′′
Vakh
750 90
298
33
RJ-23
RF53
61∘34′27.4′′
77∘46′35.4′′
Mokhovaya
1260
192.3
34
RJ-17
BL-29; RA-20
RF57
61∘11′52.7′′
75∘25′20.2′′
Vatinsky Egan
3190
287
35
BL-28
61∘12′19.5′′
75∘23′06.5′′
Er-Yakh
9.35
173
36
RJ-18
RA-19; BL-27
61∘19′41.2′′
75∘04′0.3′′
Ur'evskii Egan
359
272
37
RJ-19; R-9
BL-26
RF56
61∘26′13.6′′
74∘47′39.7′′
Agan
27 600
291
38
RJ-20
RA-18
RF55
61∘27′17.3′′
74∘40′23.3′′
Kottym'egan
7.18
192
40
RJ-21
RF54
61∘29′46.6′′
74∘15′30.3′′
Segut-Yagun
3.37
192
41
RJ-22
RF13
61∘29′11.1′′
74∘09′42.9′′
Vach-Yagun
1.79
192
42
RJ-24
RF52
61∘50′28.6′′
70∘50′28.2′′
Vachinguriyagun
9.52
192
43
RJ-25
RF14
61∘58′05.1′′
73∘47′03.4′′
Lyukh-Yagun
21.6
192
44
RF51
61∘59′39′′
73∘47′39′′
Limpas
1648
320
45
RJ-26; R-7; R-8
RA-17
RF50
62∘07′50.0′′
73∘44′05.6′′
Òromyegan
10 770
263
46
RJ-57
RF49
62∘33′39.8′′
74∘00′29.5′′
Pintyr'yagun
33.5
192
47
RJ-56
BL-25
RF48
62∘37′08.4′′
74∘10′15.9′′
Petriyagun
9.65
192
48
RJ-54; R-6
BL-24
RF47
63∘38′23.4′′
74∘10′52′′
Kirill-Vys'yagun
598
225
49
RJ-55
BL-23
RF46
62∘43′09.9′′
74∘13′45.9′′
Ai-Kirill-Vys'yagun
24.0
192
50
BL-22
RF45
63∘11′19.3′′
74∘36′25.5′′
Pyrya-Yakha
82
194
51
RA-14
63∘11′40.68′′
74∘38′16.92′′
Itu-Yakha
250
194
52
RA-13
63∘10′3.48′′
74∘45′16.32′′
Nekhtyn-Pryn
96
194
53
RA-4
63∘10′4.68′′
76∘28′19.08′′
Nyudya-Pidya-Yakha
79.5
194
54
RA-12
63∘9′31.38′′
75∘3′2.58′′
Ponto-Yakha
19
194
55
RA-11
63∘9′39.84′′
75∘09′10.86′′
Velykh-Pelykh-Yakha
170
194
56
RA-10
63∘13′12.06′′
75∘38′52.26′′
Yangayakha
88
194
57
RA-9
63∘13′25.2′′
76∘5′23.04′′
Tlyatsayakha
43
194
58
RA-8
63∘13′3.66′′
76∘15′24.6′′
Chukusamal
121
194
59
RA-3; RA-7
63∘46′22.92′′
76∘25′28.86′′
Vyngapur
1979
324
60
RA-6
63∘12′43.38′′
76∘21′27.66′′
Goensapur
11
194
61
RA-5
63∘12′45.96′′
76∘24′1.32′′
Denna
15
194
62
RA-15
63∘8′34.02′′
74∘54′29.1′′
Nyudya-Itu-Yakha
32
194
63
RJ-53; R-5
RA-16; BL-21
RF38
63∘22′01.6′′
74∘31′53.2′′
Kamgayakha
175
194
64
RJ-52; R-4
BL-19
RF39
63∘36′48.2′′
74∘35′28.6′′
Khatytayakha
34.6
194
65
RJ-51
BL-18
RF40
63∘40′41.8′′
74∘35′20.7′′
Pulpuyakha
281
194
66
RJ-50; R-3
BL-17
RF41
63∘49′58.0′′
74∘39′02.5′′
Khanupiyakha
74
194
67
RJ-29; R-2
BL-16
RF42; RF37
63∘51′23.4′′
75∘08′05.6′′
Kharucheiyakha
820
292
68
R-1; Z-55; RJ-28
BL-20; RA-2; BL-15
RF43
63∘49′54.2′′
75∘22′47.1′′
Pyakupur
9880
324
69
RJ-27; Z-86
BL-14; RA-1
RF44
63∘47′04.5′′
75∘37′06.8′′
Lymbyd'yakha
115
194
70
BL-13
63∘43′37.9′′
75∘59′04.1′′
Chuchi-Yakha
1396
292
71
RJ-32
64∘12′08.4′′
75∘24′28.4′′
Ngarka-Tyde-Yakha
59.9
186
72
RJ-30
RF36
64∘06′50.7′′
75∘14′17.3′′
Ngarka-Varka-Yakha
67.1
186
73
RJ-31
64∘09′06.4′′
75∘22′18.1′′
Apoku-Yakha
18.8
186
74
RJ-33
RY 14-49
RF35
64∘17′31.9′′
75∘44′33.4′′
Etu-Yakha
71.6
186
75
RJ-34
64∘19′10.1′′
76∘08′26.7′′
Varka-Yakha
105
186
Continued.
Number on the map
Season
N
E
River
Watersheds,
Annual runoff,
km2
mm yr-1
June
August
October
February
76
RY 14-48
64∘23′30.6′′
76∘19′50.1′′
Khaloku-Yakha
53
186
77
RJ-35
RF34
64∘26′05.2′′
76∘24′37.0′′
Kharv'-Yakha
46.4
186
78
RJ-36
RY 14-47
RF33
64∘32′07.9′′
76∘54′21.3′′
Seryareyakha
15.2
186
79
RJ-37
RY 14-46
RF32
64∘40′14.0′′
77∘05′27.2′′
Purpe
5110
309
80
RJ-38
64∘55′55.1′′
77∘56′08.2 ′′
Aivasedapur
26 100
309
81
RJ-39
RF31
65∘06′48.8′′
77∘47′58.8′′
Tydylyakha
7.46
185
82
RJ-40
RY 14-45
RF30
65∘12′17.6′′
77∘43′49.8′′
Tydyotta
12.0
309
83
RJ-41
RY 14-44
RF29
65∘23′34.1′′
77∘45′46.7′′
Ponie-Yakha
78.9
185
84
RJ-42
RY 14-43
RF28
65∘41′51.1′′
78∘01′05.0′′
Yamsovey
4030
309
85
RY 14-42
65∘46′34.5′′
78∘08′25.8′′
Khiroyakha
183
185
86
RJ-43
RF27
65∘47′48.6′′
78∘10′09.0′′
Almayakha
106
185
87
RJ-45
RF25
65∘58′54′′
77∘34′05′′
Yude-Yakha
42.4
185
88
RJ-46
RF26
65∘59′05.7′′
77∘40′52.6′′
Tadym-Yakha
39.9
185
89
RJ-44
RY 14-41
65∘57′05.5′′
78∘18′59.1′′
Pur
112 000
298
90
RJ-49
RT2 14-32
65∘59′14.7′′
78∘32′25.2′′
Malaya Khadyr-Yakha
512
278
91
RJ-48
RT2 14-31
66∘17′10.8′′
79∘15′06.1′′
Ngarka Khadyta-Yakha
1970
277
92
RT2 14-30
66∘59′20,9′′
79∘22′30.5′′
Malokha Yakha
157
208
93
RT2 14-29
67∘10′54.8′′
78∘51′04.5′′
Nuny-Yakha
656
312
94
RJ-47
RT2 14-40
67∘22′13.28′′
79∘00′25.9′′
Taz
150 000
330
95
RF21
67∘24′39′′
76∘21′12′′
Khadutte
5190
346
(a) Map of the study site with permafrost boundaries (Brown et al., 2002;
http://portal.inter-map.com (NSIDC)), runoff contour lines (Nikitin and
Zemtzov, 1986) and sampling points along the latitudinal transect of river
basin of the Ob, Pur and Taz. The numbers of the sampling sites are listed in Table
1. (b) Detailed map of the four rectangles in (a).
Chemical and isotope analyses and statistical treatment
Altogether, 95 river samples were collected in early June 2013 (spring
flood), August 2013 and 2014 (summer baseflow), October 2013 (autumn) and
February 2014 (winter baseflow) along the 1500 km latitudinal gradient
(Table 1). All sampled rivers of the WSL belong to the Kara
Sea basin. Seasonal sampling covered a full gradient from south to north,
except the month of October, which was sampled only in rivers south of
60∘ N (12 rivers in total). The watershed area of sampled rivers
ranged from 2 to 150 000 km2, not considering the Ob River in its
medium course zone. Collected water samples were immediately filtered in
pre-washed 30 mL PP Nalgene® flasks through
single-use Minisart filter units (Sartorius, acetate cellulose filter) with a
diameter of 25 mm and a pore size of 0.45 µm. The first 20 to
50 mL of filtrate was discarded. Filtered solutions for cation analyses
were acidified (pH ∼ 2) with ultrapure double-distilled HNO3 and
stored in pre-washed HDPE bottles. The preparation of bottles for sample
storage was performed in a clean bench room (ISO A 10 000). Filtered samples
for DOC, DIC, UV280nm absorbance and anions were stored in the
refrigerator for a maximum of 3 weeks before the analyses. The effect of storage for
DOC, DIC and optical measurements in boreal waters was found to be within the
uncertainty of analysis (Ilina et al., 2014). Blanks were performed to
control the level of pollution induced by sampling and filtration. The DOC
blanks of filtrate never exceeded 0.1 mg L-1, which is quite low for
the organic-rich river waters sampled in this study (i.e.,
10–60 mg L-1 DOC). pH was measured in the field using a combined
electrode calibrated against NIST buffer solutions (pH of 4.00 and 6.86 at
25 ∘C). The accuracy of pH measurements was ±0.02 pH units. DOC
and DIC were analyzed using a carbon total analyzer (Shimadzu TOC VSCN) with
an uncertainty better than 3 %. Special calibration of the instrument for
analysis of both form of dissolved carbon in organic-rich, DIC-poor waters
was performed as described elsewhere (Prokushkin et al., 2011). Major anion
(Cl, SO4) concentrations were measured by ion chromatography (HPLC,
Dionex ICS 2000) with an uncertainty of 2 %. The UV absorbance at 280 nm
is used as a proxy for aromatic C and source of DOM in the river water. It
was measured using a 1 cm quartz cuvette in a CARY-50 UV–VIS spectrophotometer
(Bruker, UK). Major cations (Ca, Mg, Na, K) and Si were determined with an ICP-MS Agilent ce 7500 with In and Re as internal standards and three various external standards, placed once per 10 samples of river water.
Approximately 30 % of samples were analyzed for Ca, Mg and Na
concentration using atomic absorption spectroscopy (flame) with an
uncertainty of 2 %. Reasonable and non-systematic agreement (between 5
and 10 %) with the results of ICP MS analyses was achieved. Aqueous Si
concentrations were also determined colorimetrically (molybdate blue method)
with an uncertainty of 1 % using a Technicon automated analyzer. The
SLRS-5 (riverine water reference material for trace metals certified by the
National Research Council of Canada) was used to check the accuracy and
reproducibility of each analysis (Yeghicheyan et al., 2014).
The 13C in dissolved inorganic carbon was analyzed in filtered river
water sampled in bubble-free sealed glass bottles by gas chromatography and
isotope mass spectrometry, using Delta V Advantage and Finnigan GasBench II
in order to determine δ13CDIC (per mil relative to V-PDB;
Fritz and Fontes, 1980). For these measurements, 0.1 mg of 100 %
H3PO4 was added to the borosilicate vial and flushed with He
(purity of 7.0) for 400 s. Afterwards, 1 mL of the sample was injected into
the vial and shaken for 36 h at 24 ∘C. Standard samples of C-O-1
and NBS-19 were routinely analyzed to test the accuracy of our measurements;
typically, a disagreement of less than 0.3 ‰ between the measured
and certified values was observed, with a total estimated measurement
uncertainty of ±0.2 ‰.
The concentration of carbon and major elements in rivers were treated using
the least-squares method, Pearson correlation and one-way ANOVA (SigmaPlot
version 11.0, Systat Software, Inc.). The ANOVA was used to reveal the
differences between different permafrost zones. It was carried out using
Dunn's method because each sampling period contained a different number of
rivers. Regressions and power functions were used to examine the
relationships between the dissolved component concentrations and the
watershed area, river discharge, average latitude of the watershed and
seasons. Comparison of DOC and major element concentration in rivers sampled
in three main permafrost zones (continuous, discontinuous and permafrost-free
regions), during all seasons and of different watershed size class, was
conducted using the non-parametric H-criterion Kruskal–Wallis test. First, we
separated the watershed into four main classes encompassing all studied
rivers (except the Ob): < 100, 100 to 1000, 1000 to 10 000, and
> 10 000 km2 . We considered three main seasons in six different
ranges of latitude (56 to 58∘ N, 58 to 60∘ N, 60 to
62∘ N, 62 to 64∘ N, 64 to 66∘ N and 66 to
68∘ N). We checked for the variation in measured parameters of each
watershed size as a function of latitude, separately in each season. In
addition, a generalized assessment of the role of permafrost type and
abundance on river water chemical composition was possible via separating all
the sampled watersheds into three categories according to the permafrost
distribution in the WSL: permafrost-free, discontinuous and continuous
permafrost.
Results
Results of major element analysis in rivers are listed in Table S1 of the
Supplement and the main results of statistical treatment are listed in Table
S2. Based on the Kruskal–Wallis H statistics, the differences between the
seasons and between different latitudes were found to be significantly higher
(p level < 0.0001) for most elements than the difference between
watershed size classes, within each season and within each latitude range.
This is illustrated for pH, DOC, DIC and δ13CDIC in
Figs. 2, 3, 4 and 5, respectively, which show the measured value as a
function of latitude for different watershed classes, individually for each
main season. Similar plots for major cations (Ca, Mg, K) and Si are given in
Supplement Figs. S1, S2, S3 and S4, respectively. The latitudinal coverage
of October was too small to be presented in these figures; however, the
October data of 12 rivers were used for statistical treatment and for
assessing the permafrost impact. There is a clear and significant trend of
concentration with latitude; the differences between different latitude
ranges are significant at p< 0.0001 for all elements, and at p< 0.05
for Si. The effect of the watershed size on river water chemical composition
in summer, winter and spring is much smaller than that of latitude
(9 < H < 12, p < 0.05 and 20 < H < 50, p< 0.001,
respectively). Considering all rivers simultaneously, the effect of the
season is clearly seen at p< 0.001 for all elements except DOC; the
latter, however, is also statistically significant (H=10.6, p=0.014).
Considering the full data set of all seasons and watershed sizes, we
distinguished three geographical zones in terms of the permafrost abundance:
continuous, discontinuous and absent. For most river water parameters (pH,
DIC, DOC, major anions and cations) the differences between three zones are
significant (30 < H < 95, p level < 0.001). Si concentration
exhibited lower but statistically significant differences between different
zones (H=9.5, p=0.0086).
Decrease in river water pH with the increase in the
latitude during winter (a), spring (b) and summer (c). The spring acid pulse
is seen only in permafrost-affected rivers north of 62∘ N (b), and
the scatter of the values is maximal during summer (c). The variability among
different watershed sizes is smaller than that between the seasons and within
the latitude gradient. Diamonds, squares, triangles and circles represent
watersheds of size < 100, 100 to 1000, 1000 to 10 000, and > 10 000 km2, respectively.
Decrease in DOC with latitude during winter (a), spring
(b) and summer (c). The latitudinal trend is significant at p< 0.05.
Considering all seasons together, the differences between different watershed
sizes are not statistically significant (p> 0.05). The symbols are the
same as in Fig. 2.
Significant decrease in DIC with latitude during winter (a), spring (b) and summer (c). Note the logarithmic scale on concentration
in all three plots. The symbols represent different size of the watershed;
see Fig. 2.
The variation in δ13CDIC with latitude
during winter (a) and spring (b) for watershed of different size. The symbols
are the same as in Fig. 2. Isotopically light DIC is observed in
permafrost-affected zone during spring, suggesting intensive respiration of
soil or plant litter carbon (Ob River sediments are from -25 to
-27 ‰; Guo et al., 2004a).
Considering all seasons and watershed sizes revealed a significant decrease
in pH, Ca and Mg northward with the largest changes occurring at the
beginning of discontinuous permafrost coverage (Fig. S5a, b and c in the Supplement,
respectively). The DOC and DIC also decrease in concentration with the
increase in the degree of permafrost coverage (Fig. S6a and b, respectively),
whereas the isotopic composition of the DIC becomes progressively more
negative northward (from ca. -15 ‰ in the permafrost-free zone to
-20 to -25 ‰ in the continuous permafrost zone, Fig. S6c). In
contrast, the effect of the permafrost on Si concentration is not clearly
seen; the scatter of the data between different seasons and watersheds does
not allow for any significant trend to be traced (not shown).
The optical properties of DOC remain essentially constant throughout the full
range of watershed sizes, latitudes and seasons (Fig. S7). The largest
variation in specific UV280nm absorbance occurred in winter, when
several DOC-rich waters from the southern (permafrost-free) part of the WSL
demonstrated quite a low concentration of aromatic (colored) compounds.
Discussion
Effect of latitude (permafrost and vegetation) on major cation, Si and DIC mobilization from the soil profile and groundwater to the river
From general knowledge of environmental control on carbon and major element
fluxes in rivers of the Russian subarctic (Prokushkin et al., 2011; Pokrovsky
et al., 2012) and other boreal and subarctic regions (Laudon et al., 2004;
Petrone et al., 2006; Walvoord and Striegl, 2007; Jantze et al., 2013;
Giesler et al., 2014), we anticipate a decrease in most element
concentrations, including DOC, northward regardless of the season and the
river size in the WSL due to (1) a decrease in chemical weathering
intensity with the temperature, well demonstrated for igneous rocks such as
basalts (Dessert et al., 2003) and granites (Oliva et al., 2003); (2) a
decrease in the thickness of peat deposits in total and the active soil
(peat) layer in particular (Liss et al., 2001; Beilman et al., 2009;
Stepanova et al., 2015, and references therein); (3) a decrease in plant
biomass and related plant litter stock on the surface of the soils (Tyrtikov,
1979; Frey and Smith, 2007); (4) a shortening of the unfrozen period of the
year; and (5) a decrease in the degree of groundwater feeding (Romanovsky,
1983; Nikitin and Zemtsov, 1986; Fotiev, 1991). The factors capable of
enhancing element export fluxes in northern (permafrost-bearing) rivers
relative to southern (permafrost-free) rivers of the WSL are (1) the
decrease in dissolved organic matter (DOM) respiration by heterotrophs in the
water and soil column and thus the increasing removal of allochthonous DOC
from the soil to the river (Striegl et al., 2005); (2) the increase in DOC
and related element leaching from plant litter and topsoil (Pokrovsky et al.,
2005; Giesler et al., 2006; Fraysse et al., 2010) during more pronounced
massive freshet event or summer high flow (Michel and Vaneverdingen, 1994;
McClelland et al., 2006; White et al., 2007); (3) the decrease in DOM
retention (adsorption) on the mineral soil horizon because clay horizon is
typically frozen in the north (Kawahigashi et al., 2004); (4) the decrease
in authigenic clay and allophane mineral formation in the soil horizons
(Targulian, 1971).
At the current, rather limited, stage of knowledge of mineral, organic soil
horizons and plant biomass chemical composition and reactivity across the
WSL, only a few environmental factors can be quantitatively tested based on
river water chemical analyses. In the case of the dominance of groundwater
feeding of the river, the decrease in element concentrations from water–rock
interaction whose transport is not limited by availability of DOM (Ca, Mg,
DIC) is expected to be most pronounced in winter, when the groundwater
feeding is maximal (see Walvoord and Striegl, 2007, for the Yukon River basin
example). Moreover, in the permafrost-bearing zone during winter baseflow,
one should expect significant differences in element concentration in winter
between small rivers (weakly or not affected by taliks) and large rivers
(essentially fed by taliks), as is known from local geocryological
conditions (Baulin et al., 1967; Romanovsky, 1983; Fotiev, 1989, 1991; Ivanov
and Beshentsev, 2005). In spring, when the active layer is very thin and the
majority of the soil column is frozen, the export from the watershed is
dominated by surface flow and thus the difference in groundwater-related
element concentration between (i) small and large rivers and (ii) north and
south should be minimal. However, the abovementioned hypotheses are not
supported by DIC, Ca and Mg concentrations observed in rivers (Figs. 4, S1
and S2). First, the DIC concentrations decrease between permafrost-free and
discontinuous/continuous permafrost zones is a factor of 15 ± 5 in
winter (Fig. 4a) and a factor of 60±10 in spring (Fig. 4b). Similarly,
the decrease in Ca and Mg concentrations between south of 59∘ N and
62–66∘ N zones is 10-fold in winter and 20–30-fold in May.
In fact, it is the spring period which exhibits the highest contrast in
element concentrations between the south and the north. Second, for the
latitude concentrations gradient from south to north, the relative DIC, Ca
and Mg concentration change between large (1000–10 000 and
> 10 000 km2 ) and small (< 100 km2) rivers in winter is
not statistically significant (p> 0.05).
However, a systematic decrease in Ca concentration in the WSL rivers
northward (Figs. S1, S5b) is consistent with a general decrease in Ca
concentration in soil ecosystems as illustrated in Fig. S8. An order-of-magnitude decrease in Ca concentration in mineral horizons of WSL peat
columns occurred between 55 and 66∘ N (Stepanova et al., 2015). On a
smaller scale, a 3-fold decrease in exchangeable Ca concentration in
alluvial soils of the Ob Basin from 56 to 60∘ N was reported
(Izerskaia et al.,
2014). These observations confirm a strong control of lithology and soil
weathering on Ca concentration in both deep and surface soil horizons and
vegetation, which finally determines the extent of Ca transport via surface
flux to the river.
North of 66∘ N, concentrations of Ca, Mg and sulfate increase
relative to their concentration at 62–66∘ N of discontinuous
permafrost zone. This is especially pronounced during the summer period
(Figs. S1c, S2c). We do not exclude here the influence of marine sedimentary
deposits containing salts in the deep part of the mineral soil profile below
the peat layer. These deposits are described in the low reaches of Taz and
Pur rivers, based on sedimentary cores extracted during extensive drilling of
the territory (Liss et al., 2001). This influence, however, cannot be
unequivocally evidenced because (i) DIC concentrations also increase in
summer, north of 66∘ N, although DIC is not likely to be affected by
marine deposits, and (ii) chloride, an efficient marker of sea salts, is
not increasing in the north (not shown).
The isotopic composition of DIC confirms the general features of DIC and
cation concentration (Fig. 5). The groundwater feeding by taliks in winter is
highly uniform over 10∘ of latitude, with the value of δ13CDIC being equal to -15 ± 5 ‰, reflecting
both carbonate/silicate weathering and a buildup of CO2 with a stronger
respiratory signal (Finlay, 2003; Striegl et al., 2001; Giesler et al., 2014;
Rinta et al., 2015). During this period, the variability in
δ13CDIC is the highest in small (< 100 km2)
watersheds, but no trend of isotopic composition with latitude could be
evidenced at p< 0.05 (Fig. 5a). This isotopic signature is preserved in
spring for southern (< 60∘ N) watersheds whereas in
permafrost-affected regions, δ13CDIC decreases to ca.
-25 to -20 ‰ regardless of the river size and the type and
the abundance of the permafrost (Fig. 5b). Such low values in the
permafrost-affected zone could no longer represent the influence of
carbonate/silicate rock weathering by soil CO2 and likely reflect direct
microbial processing of soil and sedimentary organic matter (Waldron et al.,
2007; Giesler et al., 2013), with the DIC isotopic signature similar to that
of organic carbon in western Siberian subarctic topsoil
(-26 ± 2 ‰; Gentsch et al., 2015) and the Ob River organic
sediments (-25 to -27 ‰; Guo et al., 2004a).
A plausible explanation for the δ13CDIC seasonal
variation being mostly pronounced in the permafrost zone can be that
microbial mineralization of dissolved organic carbon occurs most efficiently
during the springtime, when significant amounts of fresh organic matter from
ground vegetation are leached by melted snow. Higher bioavailability of
vegetation leachates relative to more refractory soil humic and fulvic acids
is known from studies in other temperate (van Hees et al., 2005) and boreal
(Wickland et al., 2007) regions. The lack of δ13CDIC
decrease in spring relative to winter in the permafrost-free zone may stem
from (i) significant input of the carbonate/silicate rock-hosted groundwaters
during the full period of the year in the south or (ii) the different nature of
DOM in the south, where the more refractory organic matter originated from
peat leaching is less subjected to microbial processing compared to fresh
vegetation leachates in the north, where the peat soil in spring is frozen.
One has also take into account that the DIC concentrations in spring are a
factor of 30 lower in the permafrost-bearing region relative to the
permafrost-free region (Fig. 4b). As such, a relatively small input of
microbially respired CO2 will be significantly more visible in the
δ13CDIC value of the northern rivers compared to that of
the southern rivers.
The variation in δ13CDIC along the
permafrost and latitude gradient helps to better explain the origin of DIC in rivers in contrasting
permafrost zones. Consistent with a progressive decrease in the groundwater
feeding of rivers northward (Nikitin and Zemtzov, 1986; Frey et al., 2007b),
we observe a distinct trend of δ13CDIC with the latitude
during the spring period, reflecting the shift of DIC origin from groundwater in
the south to plant litter degradation and soil respiration in surface waters
north of 62∘ N (Fig. 5b). In winter, the δ13CDIC is rather constant within the full latitudinal profile
(Fig. 5a), confirming the dominant role of carbonate/silicate mineral
weathering by atmospheric and soil CO2 dissolved in the groundwaters.
Indeed, hydrological studies in the WSL revealed that the groundwater feeding
of small (< 10 000 km2 watershed) rivers decreases from
20–30 % in the discontinuous and sporadic/isolated part of the WSL to
3–6 % in the northern, continuous permafrost zone (Novikov et al.,
2009). These numbers agree with estimations based on Russian Hydrological Society (RHS) data of large
western Siberian rivers (Nadym, Pur and Taz) and the left tributaries of the
Yenisei River (Dubches, Elogyi and Turukhan; Nikitin and Zemtzov, 1986).
According to more recent evaluations of Frey et al. (2007b), the groundwater
contribution to summertime period river chemical composition ranges between
30 and 80 % for the rivers located between 56 and 58∘ N.
Consistent with these findings, the pH values of 7 to 7.5 in the southern
rivers observed both in winter and spring (Fig. 2a, b) are indicative of
carbonate/silicate rock input. The spring acid pulse, well established in
other permafrost-free boreal regions (Buffam et al., 2007), is not at all
pronounced in the south of the WSL but becomes clearly visible in the
permafrost-affected, northern regions where the springtime pH decreases to
5 ± 0.5 (Fig. 2b). This illustrates the more important role of plant litter
and moss leaching in the permafrost-bearing zone on solute export from the
watershed. In addition, the dominance of sands north of 62∘ N (Liss
et al., 2001) may allow low-molecular-weight (LMW) organic acids migrate to
the river from the soil profile. In the southern, permafrost-free zone, the
dominating clays underneath the peat can adsorb acidic LMW organic compounds
and thus do not allow the acid pulse to be clearly visible.
The increase in pH in summer relative to the spring period is again less visible
in the south than in the north (Fig. 2c) and may reflect the persisting role
of bedrock dissolution as well as the change in the river feeding regime,
from top soil and vegetation in the north to the peat soil column leaching in
the south. The summertime increase in river water pH north of 60∘ N,
in the forest–tundra and tundra zone may be linked to (i) enhanced
photosynthesis in rivers of the north due to better insolation and less
forest shading and (ii) mobilization of DOM and other solutes from soil
depressions rather than from watershed divides. The depressions are subjected
to intense rinsing during the spring seasons, when the majority of soluble
acidic compounds are flushed from the litter and Oe horizon.
These mechanisms are evidenced from studies of the hydrological balance of
frozen bogs performed in the northern part of studied territory (Novikov et
al., 2009). In contrast, the watershed divides contain significant amounts of
organic litter and release organic acids only in spring, when they are
covered by temporary ponds of melted snow (see Manasypov et al., 2015). This
hydrological scheme of river water feeding is based on the seasonal
multiannual observations on frozen bogs of the north of the WSL (Novikov et al.,
2009), although the chemical nature of DOM mobilized from different parts of
the watershed remains unknown.
The importance of plant litter and ground vegetation leaching as element
sources in western Siberian rivers can be assessed from the comparison of K
concentrations as a function of latitude during different seasons (Fig. S3).
The most significant decrease in K concentration from the southern
(< 59∘ N) to the northern (61–67∘ N) watersheds occurs
in spring, during intense plant litter leaching. Regardless of latitude, K
concentration follows the order spring > winter > summer, with the
highest concentrations, up to 2500 ppb, recorded in permafrost-free region.
Given that the other cations, possibly originating from the water–mineral
interaction at some depth, do not exhibit such high concentration in spring,
we interpret the springtime K “pulse” as indicative of plant litter
leaching in the productive taiga zone. This “pulse” is much less visible in the
permafrost zone due to significantly lower biomass and primary productivity
of forest–tundra and tundra biomes compared to the taiga of the WSL
(Tyrtikov, 1979; Liss et al., 2001).
Despite significant variability in Si concentrations among rivers of various
sizes across the latitude profile (Fig. S4), the concentrations in the
permafrost zone are not lower than those in the south of the WSL. Results of
a previous study of WSL rivers during summer show that Si concentrations are
weakly dependent on latitude (Frey et al., 2007), as also confirmed in this
work for the spring flood and winter baseflow period. Given that (i) the
dominance of permafrost north of 64∘ N implies very low groundwater
feeding (4 to 6 % of the annual discharge; see Nikitin and Zemtsov, 1986;
Novikov, 2009) and (ii) the upper part of the soil profile including its
seasonally frozen and unfrozen parts is mostly peat rather than silicate
mineral sediments, the role of groundwater–silicate rock interaction in Si
supply to northern rivers should be quite low. Therefore, we hypothesize that
elevated concentrations of Si in northern rivers are due to peat leaching and
degradation. A depletion of Si in rivers of the southern part of the WSL may
be due to Si retained by abundant bog and forest vegetation. This is
consistent with the general setting of the WSL, recovering from the last
glaciation (Liss et al., 2001), with contemporary peat accumulation in the
south and old frozen peat thawing/degrading in the north.
DOC concentration across a 1500 km latitude transect of variable permafrost coverage
Results of organic carbon concentration in western Siberian rivers collected
over various seasons of the year generally confirm the pioneering findings of
Frey and Smith (2005). The strong statistically significant (p< 0.05)
contrast in DOC concentration between permafrost-free, discontinuous and
continuous permafrost zone persists over the course of the year and each
season except probably winter (Figs. 3 and S6a). This difference is also seen
in δ13CDIC values among all three zones (Fig. S6c),
suggesting, on the annual scale, a more significant contribution of microbial
processing of plant and soil organic carbon to HCO3 and CO2 of the
river water in the permafrost-bearing zone compared to the permafrost-free
zone.
In accordance with the conclusion reached by Frey and Smith (2005), the
variation in hydrology may play a limited role in DOC variability and export
from the watershed of WSL rivers. The gradient in DOC concentrations along
the latitudal profile remains similar between spring flood and summer
baseflow (Fig. 3b and c). Although the winter period does not exhibit such a
clear difference between permafrost-free and permafrost-affected regions
(Fig. 3a), the contribution of the winter discharge to the annual flux of DOC
is between 10 and 15 % and as such does not significantly affect annual
export of DOC from the watersheds.
In contrast, the gradient of organic carbon concentration along the
latitudinal profile in spring will be mostly controlled by the difference in
plant litter stock subjected to leaching by melted snow. As such, one would
not expect any significant difference between large and small rivers at
otherwise similar runoff, vegetation and bog coverage. This is partially
confirmed by the similarity of the UV280nm–DOC slope,
corresponding to similar degree of DOM humification, among different seasons
and latitudinal positions (Fig. S7). The uniform distribution of UV280
absorbance demonstrates that the main control of DOC by allochthonous (terrestrial)
input from peat and/or ground vegetation leachates. The exceptions are the
rivers Vasyugan (no. 21), Shegarka (no. 4) and Vatinsky Egan (no. 34),
exhibiting low UV280nm at high [DOC] (Fig. S7). These rivers are
potentially affected by oil production sites and may contain some uncolored
products of hydrocarbon oxidation in the underground waters.
Overall, results on western Siberian rivers generally confirm the conclusion
of Finlay et al. (2006) on (i) the lack of groundwater contribution to
streamflow in arctic watersheds and (ii) that river DOC dynamics are driven
essentially by processes occurring at the soil surface. However, we doubt the
importance of large DOC pool production under very cold conditions with regard to the
main reason for sustained high concentration of DOC at snowmelt suggested by
Finlay et al. (2006). Indeed, the plant litter degradation in winter, even in
the warmest scenario, is minimal and does not contribute significantly to
annual litter leaching (Bokhorst et al., 2010, 2013). Instead, we suggest
fast plant litter and ground vegetation leaching in spring, at the very
beginning of the snow melt. Such a fast enrichment in DOC and colored organic
compounds of surface water depressions, on the order of several hours, has
been observed in the discontinuous permafrost zone in early June (Manasypov
et al., 2015). Significant release of DOC and nutrients from flooded ground
vegetation in the southern part of the WSL is also known (Vorovyev et al., 2015).
An unexpected result of the study of western Siberian watersheds is the lack
of the enrichment in DOC of small headwater streams, in contrast to what has
been reported for Scandinavian rivers and streams (Ågren et al., 2007, 2014,
and references therein). In the WSL, especially in the northern,
permafrost-affected zone, the small (< 100 km2) streams yielded DOC
concentrations that were not statistically higher (p> 0.05) than those of
larger rivers, neither in spring flood nor in summer. A number of factors can
be responsible for the observed difference between permafrost-free European
and permafrost-bearing Siberian watersheds. In the north of western Siberia,
the microbial processing of DOM in large rivers may be weakly pronounced.
This is confirmed by the observation that the degree of light C isotope
enrichment (lowering δ13CDIC) in spring is independent
(p> 0.05) of the size of the river (Fig. 5b) and, correspondingly, of
the water residence time on the watershed. According to Kawahigashi et
al. (2004), the DOM in northern, permafrost-affected tributaries of the
Yenisey River was significantly less biodegradable than that in southern
tributaries. This may contribute to better preservation of DOM in the stream
yielding its independence of the water travel time. Small watersheds of
western Siberia exhibit a runoff and average slope very similar to that of
the large rivers, given the very flat orographic context of the WSL. This
contrasts with the mountain regions of Sweden and Alaska, where the headwater streams may
exhibit higher runoff and thus higher export of the dissolved constituents.
Finally, the riparian zone, very important for regulation DOC stock and
export in small streams draining glacially formed terrain of NW Europe (Dick
et al., 2015; Kuglerová, et al., 2014), is much less pronounced in
western Siberia, where generally flat, frequently flooded areas dominate the
watershed profile.
The elevated DOC concentrations in continuous permafrost zone, especially
north of 67∘ N observed in the present study (Fig. 3b, c), are
consistent with previous results showing that, for otherwise similar factors,
the permafrost areas are a greater source of DOC than the areas with seasonal
frost (Carey et al., 2003). In permafrost areas, meltwater travels through
organic-rich layers in the form of so-called supra-permafrost flow, as
opposed to areas without impermeable permafrost table. In the latter, the
infiltration of organic-rich surface waters to the deep mineral layer and DOC
sorption on clay minerals may occur, thus decreasing the overall export of DOC
(see Smedberg et al., 2006, for discussion). Given the dominance of peat
rather than minerals within the active (unfrozen) layers of soil profile, the
difference between permafrost-free and permafrost-affected zones is even more
accentuated in western Siberia.
A sketch of typical soil profiles of western Siberia in the permafrost-free
and permafrost-bearing zone presenting DOC mobilization pathways from the
soil to the river in the end of active period is shown in Fig. 6. The
two cross sections shown in this figure are highly representative for two
most contrasting cases of soil and watershed flux formation, corresponding to
dark coniferous taiga in the permafrost-free zone and dwarf shrubs with green
mosses of tundra and forest–tundra in frozen peatlands of continuous
permafrost zone; both sites are located at the watershed divide. The detailed
position of soil horizons and their attribution to FAO is based on available
literature data (Tyrtikov, 1973, 1979; Liss et al., 2001; Pavlov and
Moskalenko, 2002) and our recent investigations of the region (Loiko et al.,
2015; Stepanova et al., 2015). We hypothesize that plant-litter- and topsoil-derived DOC adsorbs on clay mineral horizons in the southern, permafrost-free
and discontinuous/sporadic permafrost zone but lacks the interaction with
minerals in the continuous permafrost zone. This assumption corroborates
results found during another latitudinal river transect of Siberia, along the
Yenisey River and its left tributaries draining peatlands of the WSL
(Kawahigashi et al., 2004, 2006): the northern tributaries exhibited
significantly higher DOC concentrations than the southern tributaries of this
river. Specifically, given the significant thickness of the peat even in the
northernmost part of the WSL and the active layer thickness of < 50 to
80 cm (30 cm on mounds and 80 to 150 m in troughs and depressions,
Tyrtikov 1973, 1979; Baulin et al., 1967; Baulin, 1985; Khrenov, 2011;
Novikov et al., 2009), even in the region of continuous permafrost
development, peat soil interstitial solutions might not come in contact with
the mineral soil horizon and thus will not decrease their DOC concentration
during migration from the soil column to the river along the permafrost
impermeable layer (Fig. 6).
Scheme of DOC pathways within the soil profile and to the river.
(a) In forest watershed of the south, permafrost-free zone
(57∘ N). Soil horizons (FAO, 2006): (1) O (Mor, forest litter),
(2) medium-decomposed peat (He) transforming into strongly decomposed peat (Ha)
in the bottom layer, (3) mollic humic horizon, (4) ABg surface horizons
with stagnic properties, (5) Bg middle stagnic horizon, and (6) Cgk
carbonate-bearing clays and clay loam.
(b) DOC pathways in frozen bog peatlands of
continuous permafrost (67∘ N). Soil horizons (FAO, 2006): (7) weakly decomposed peat (Hi), (8) partially
decomposed peat (He), (9) humic horizons (AH), (10) cryoturbated frozen
stagnic horizon (Bgf), (11) frozen stagnic horizon (BCgf), and (12) sedimentary
deposits (Cf).
In the south, DOC is retained by clay, and deep in the soil profile by clay
loam with carbonates. In the north, the active layer depth does not exceed
the overall thickness of the peat and thus the leachates of ground vegetation
and peat layer do not meet mineral horizons during their transit to the
river.
Therefore, in the permafrost zone, the DOC export is strongly controlled by
DOC residence time and the water travel pathway through organic topsoil and
lichen, moss and litter leaching vs. peat and mineral layer leaching
(Fig. 6b). In this case, it is only the thickness of the unfrozen peat and
the local permafrost coverage that control the DOC export from the soil to the
river. As a result, DOC concentration in the streams will be weakly dependent
on the watershed size and seasons. It follows that DOC export from peat soils
by medium-sized (n×10 000–n×100 000 km2) rivers
located entirely in the permafrost zone may be higher than that of the larger
rivers, crossing permafrost-free regions. This hypothesis is supported by
available information on DOC yield by rivers of the WSL. Thus, the Taz (s=150 000 km2), Pur (112 000 km2) and Nadym rivers
(64 000 km2), entirely located in the discontinuous permafrost zone,
exhibit 1.9, 2.1, and 4.4 t km-2 yr-1 DOC yield, respectively
(Gordeev et al., 1996, and calculated based on data of the RHS). This is
significantly higher that the value suggested for the Ob River
(1.2 t km-2 yr-1; Gordeev et al., 1996).
Possible evolution of chemical composition and fluxes of western Siberian rivers under climate change scenarios
The most likely scenario of the climate change in western Siberia consists of
shifting the permafrost boundary further north and increase in the active
layer thickness (Pavlov and Moskalenko, 2002; Frey, 2003; Romanovsky et al.,
2010; Vasiliev et al., 2011; Anisimov et al., 2013). The permafrost boundary
change, equivalent to the northward shift of the river latitudes, may
decrease the DOC concentrations of the most northern rivers by a maximum of
2-fold due to the change of continuous to discontinuous permafrost. The thickness of
the active layer is projected to increase by more than 30 % during this
century across the tundra area in the Northern Hemisphere (Anisimov et al.,
2002; Stendel and Christensen, 2002; Dankers et al., 2011). In the WSL, this
increase will be most dramatic in the north, where the peat deposits are
thinner than those in the discontinuous permafrost zone (Botch et al., 1995;
Liss et al., 2001; Novikov et al., 2009; Kremenetski et al., 2004). Assuming a
short-term (hundreds of years) scenario in the WSL, we hypothesize that the main
consequences of this increase will be the involvement of upper clay horizon
and sand/silts in water pathways within the soil profile. As a result, the
DOC originating from the upper peat layer leaching and plant litter
degradation will be retained on mineral surfaces and in the clay interlayers
(Kaiser et al., 2007; Oosterwoud et al., 2010; Mergelov and Targulian, 2011;
Gentsch et al., 2015). To which degree this change of water pathways in the
soil column may affect the other dissolved components cannot be predicted.
However, this effect for inorganic solutes is expected to be lower than that
for DOC, given much lower affinity of HCO3, major cations and Si to clay
surfaces and the lack of unweathered (primary) silicate rocks underneath the peat
soil column. Nevertheless, the possibility of leaching of inorganic
components from the mineral layers should be considered. For example, DOC
export exceeded DIC export in a tributary of the Yukon River during high
flow, whereas DIC predominated during low flow and the DIC yields increased
with decreasing permafrost extent (Dornblaser and Striegl, 2015).
Unfortunately, no time series on hydrochemistry of rivers of continuous
permafrost development, north of 64∘ N, are available to test the
hypothesis of the impact of climate change on a possible decreasing DOC flux
from frozen peatlands and the DOC/DIC change due to ongoing decrease in
permafrost protection of the mineral layer from adsorbing DOC.
Important modifications linked to the climate change in boreal and subarctic
zones concern the change of the hydrological regime (Karlsson et al., 2015),
in particular the increase in the winter baseflow (Yang et al., 2004; Ye et
al., 2009; Serreze et al., 2000) due to the increase in the groundwater
feeding (Frey et al., 2007a, b; Walvoord and Striegl, 2007; Rowland et al.,
2010; Walvoord et al., 2012), coupled with the increase in the overall
precipitation and, consequently, water runoff (Peterson et al., 2002;
McClelland et al., 2006). Here, we argue that the 10 to 30 % modification in
the annual runoff will be within the variation in the DOC and cation
concentrations between watersheds of various sizes observed in the present
study and as such will not significantly affect the export fluxes of river
water constituents.
To which degree the ongoing DOC concentration and flux rise in rivers, linked
to climate change and/or acidification as reported in western Europe and
Canada (Worrall et al., 2004; Porcal et al., 2009) can be applied to the WSL
is unknown. However, we did not observe any significant (i.e., > 30 %)
change of DOC fluxes over past 30 to 40 years neither in the boreal
non-permafrost pristine region of NW Russia (Severnaya Dvina River; Pokrovsky
et al., 2010), nor in the Central Siberian, continuous permafrost rivers of
the Yenisei Basin (Pokrovsky et al., 2005). Moreover, a decrease in DOC
fluxes in the Yukon River was reported and suggested to be linked to enhanced
mineralization of DOC by biota (Striegl et al., 2005). Note also that the
more recent evaluation of the Ob River DOC discharge using flow-weighted
concentration of 9.4 mg L-1 measured in 2003–2007 (Cooper et al.,
2008) gives a flux of 1.3 t C km-2 yr-1, well comparable with
the earlier estimate of 1.2 t C km-2 yr-1, based on the RHS data of
1950–1990 (Gordeev et al., 1996).
The increase in vegetation productivity reported for Arctic river basins
(Sturm et al., 2001, Tape et al., 2006; Kirdyanov et al., 2012) will most
likely proportionally increase the springtime K flux due to its leaching
from plant litter but likely decrease the summertime Si flux, especially
in the permafrost-bearing regions. The increase in vegetation density in the
next decades to centuries may produce a transient uptake of Si by growing
vegetation in the discontinuous permafrost zone during summer period.
However, this potential decrease in Si export flux may be outweighed by the
increasing release of Si from previously frozen mineral horizons and as such
the overall modification of the Si concentration and riverine flux in the
discontinuous–continuous permafrost zone may be smaller than that projected
by simple latitudinal shift.