Temperature-driven nutrient recycling and euxinia as a marine mass extinction mechanism
Dominik Hülse ( [email protected] ) University of California, Riverside https://orcid.org/0000-0001-5386-6746 Kimberly Lau The Pennsylvania State University Sebastiaan van de Velde University of California, Riverside Sandra Arndt University of Bristol Katja Meyer Willamette Andy Ridgwell University of California, Riverside https://orcid.org/0000-0003-2333-0128
Article
Keywords: nutrient recycling, euxinia, temperature, microbial metabolism, ocean redox state, carbon cycling, end-Permian mass extinction
Posted Date: November 30th, 2020
DOI: https://doi.org/10.21203/rs.3.rs-80350/v1
License: This work is licensed under a Creative Commons Attribution 4.0 International License. Read Full License 1 Temperature-driven nutrient recycling and euxinia as a
2 marine mass extinction mechanism
1 2,3 1 3 Dominik Hulse,¨ Kimberly V. Lau, Sebastiaan J. van de Velde,
4 5 1 4 Sandra Arndt, Katja Meyer, Andy Ridgwell
1 5 Department of Earth and Planetary Sciences, University of California, Riverside, CA, USA
2 6 Geosciences and Earth and Environmental Systems Institute, The Pennsylvania State University,
7 University Park, PA ,USA
3 8 Geology and Geophysics, University of Wyoming, WY, USA
4 9 Bgeosys, Geoscience, Environment & Society, Universite´ Libre de Bruxelles, Brussels, Belgium
5 10 Earth and Environmental Sciences, Willamette University, Salem, OR, USA
11 Extreme warming at the end-Permian induced profound changes in marine biogeochemical
12 cycling and animal habitability, leading to the largest extinction in Earth’s history. However,
13 a causal mechanism for the extinction that explains the different proxy evidence has yet to
14 be found. By combining recent modeling developments with global and local redox obser-
15 vations, we show that a temperature-driven increase in microbial respiration can reconcile
16 reconstructions of the spatial distribution of euxinia and seafloor anoxia spanning the Per-
17 mian/Triassic transition. We illustrate how enhanced metabolic rates would have strength-
18 ened upper ocean nutrient recycling, and thus shoaled and intensified the oxygen minimum
19 zones eventually causing euxinic waters to expand onto continental shelves, poisoning ben-
20 thic habitats. Finally, we find that the temperature effect on microbial activity can account
1 21 for some of the decline in carbon isotopes at the end-Permian with the implication that car-
22 bon release as inferred from those changes is likely overestimated. Our findings present
23 a novel view of the sensitive interconnections between temperature, microbial metabolism,
24 ocean redox state and carbon cycling during the end-Permian mass extinction with potential
25 far-ranging implications for the interpretation of carbon cycle perturbations during Earth
26 history.
27 Climate warming driven by volcanic greenhouse gas release is widely regarded as the under-
28 lying driver for the largest metazoan extinction event in Earth’s history at the end of the Permian
1, 2 29 Period when ∼90% of marine species were eliminated . Although the extinction event itself has
30 been intensely studied and is relatively well characterized, the specific physical and/or biogeo-
3 31 chemical environmental changes that drove biodiversity loss in the ocean are uncertain . Proxy
◦ 4–6 32 evidence reveals a 7-10 C increase in sea surface temperature occurring in as little as ∼39 kyr
7 33 (Fig. 1a+d), the development of (photic-zone) euxinia (waters containing sulphide ), an expan-
8, 9 34 sion in the extent of seafloor anoxia , and a decrease in the carbon isotopic signature recorded in
10 13 11 35 carbonates (δ Ccarb, Fig. 1e), all approximately reaching their nadir at the extinction horizon
36 (EH, ∼251.94 Ma).
37 Proposed explanations linking these observations with the extinction all require reduced oxy-
38 genation of the ocean, but fundamentally diverge at this point. In particular, previous 3D Earth
39 system model (ESM) studies have required either a sustained collapse of global ocean circulation
12 40 in conjunction with a much weaker biological pump , or a well ventilated end-Permian ocean
2 41 in conjunction with a much stronger biological pump driven by enhanced nutrient (phosphate)
13, 14 15 42 availability . However, other modelling work has demonstrated that increasing the ocean
43 phosphate inventory sufficiently to create widespread subsurface euxinia, not only requires an ex-
16 44 cessive increase in phosphate availability, but also results in near-global anoxia in the deep sea
9,17 45 which is in conflict with paleoredox estimates from uranium isotope records (Fig. 1b+c). Other
46 explanations focus on reducing oxygen availability throughout the ocean as a whole, either through
13 47 the oxidation of methane released from hydrates or of a massive reservoir of dissolved organic
18 19, 20 48 matter , or via warming driven by the CO2 release associated with volcanism . These car-
13 49 bon release mechanisms can also drive a pronounced decline in δ Ccarb. However, a gradient in
13 21 50 δ Ccarb from surface to subsurface records suggests a vertical partitioning in ocean geochem-
51 istry occurred and that there is much more to the event than carbon release and climate warming.
52 Here we reconcile the varying proxy observations and provide new insights into the controls on
53 subsurface euxinia and hence marine extinction mechanisms, by recognising the important role of
54 temperature in stimulating microbial respiration and thus dictating the redox profile in the water
55 column.
56 Oxygen availability in the water column generally decreases from well-mixed surface waters
57 (few tens of meters), to the oxygen minimum zone (OMZ, typically at a depth of a few hundreds
58 of meters). Along with vertical ocean mixing, this gradient is controlled by the remineralization
59 of particulate organic matter (POM) which consumes oxygen and releases inorganic nutrients (and
22 60 carbon) that can be returned to the surface to fuel new primary productivity . Critical here is the
61 interplay between reactivity and sinking rate of POM, as it controls the shape of the remineraliza-
3 62 tion rate depth profile (i.e. the scale depth of POM remineralization) and, thus, the intensity and
16, 23 63 depth of OMZs . We posit that a further factor, and the key to understanding how the marine
24 64 environment changed across the Permian/Triassic boundary (P/Tr), is ocean temperature .
65 A mechanistic representation of the biological pump
66 To demonstrate the importance of a warming ocean in driving subsurface euxinia and potentially
67 widespread extinction across shallow marine environments, we simulate redox distributions for a
25, 26 68 range of P/Tr conditions using the cGENIE ESM . We modify the widely used ’static’ represen-
27 69 tation of the biological pump (i.e. an invariant POM remineralization depth profile ), to allow us
70 to explicitly account for the impact of ocean warming at the end-Permian on POM remineralization
71 (SI). In addition, to reflect the shallower remineralization profile inferred prior to the rise of pelagic
23 72 calcifiers in the early Mesozoic , we decrease the sinking rate of POM in the model by ∼22% –
28 73 scaled to the smaller mean animal biovolume at the end-Permian ( , SI). Finally, we account for
74 changes in POM remineralization as reactive POM compounds react with sulphide (H2S) and be-
29 75 come less susceptible to bacterial remineralization (SI), in a process known as “sulfurization” .
76 Because the time-scale of warming leading up to the P/Tr boundary is slow relative to the adjust-
77 ment time-scale of large-scale ocean circulation (i.e. warming likely occurred over ∼39 kyr or
6, 11 78 more, starting in the C. meishanensis biostratigraphic zone ), a series of (10 kyr) steady-state
79 simulations is appropriate for simulating the biogeochemical response to warming. In these, we
80 prescribe a range of atmospheric pCO2 (1 – 30 × pre-industrial pCO2, i.e. 280–8400 ppmv) chosen
81 to span the increase in tropical Tethys ocean temperatures reconstructed from proxy records (∼22
◦ 82 – 35 C, Fig. 1a+d, SI). Simultaneously, we explore the importance of 1 – 2.5× modern ocean
4 83 phosphate inventories to represent the potential net impact of increased weathering and sediment
30–32 84 regeneration rates as the climate warms and ocean anoxia increases, respectively – thus, creat-
85 ing model ensembles of varying climate vs. ocean nutrient state. Atmospheric oxygen is fixed at
86 modern.
87 Constraining model results with global and local redox proxies
88 Uranium isotopes can provide powerful constraints on ocean models via the reconstructed extent
89 of seafloor anoxia (fanox) . To quantitatively compare the ESM results with a compilation of car-
238 90 bonate δ U data over the P/Tr transition (SI), we use a forward box model that encapsulates the
9 91 uncertainties in the U isotope budget (adapted from Lau et al. ). According to our U-model results,
238 92 the δ U data can be best explained by an abrupt increase in fanox that either coincided with the
93 EH, or preceded it by much less than the onset of the warming event (Fig. 1 b+c). An increase
94 of fanox from a modern value of 0.6% to about 30% (i.e. a factor of 50) represents our preferred
95 scenario as other perturbations fail to simulate the rate of change and magnitude of the shift in the
238 96 δ U data (Extended Data Fig. 4).
97 We also ground-truth the ESM results with a new compilation of local redox proxies (Ex-
98 tended Data Fig. 1 and Extended Data Table 2). The data-set consists of geochemical, lithologic,
99 sedimentologic, and biomarker evidence for water-column euxinia and bottom water anoxia and
100 distinguishes three phases of the P/Tr transition (Late Permian background, Start Warming and
101 Main Extinction, Extended Data Table 2). The Late Permian data only indicate photic-zone eu-
102 xinia occurring at the Meishan section with no clear evidence for seafloor anoxia (Fig. 3a+g,
5 103 Extended Data Table 2). Immediately before the EH, water-column euxinia expanded to sections
104 in British Columbia (BC), Shangsi, and potentially to the equatorial Panthalassic (Japan), while
105 unequivocal evidence for bottom water anoxia remains limited and only exists for BC (Fig. 3b+h).
106 In contrast, during the Main Extinction, anoxia spread across much of the shallow seafloor, and
107 multiple lines of evidence suggest widespread euxinia impinged on shallow marine habitats (Fig.
108 3c+i).
109 Impact of temperature-driven respiration on the marine redox-landscape
110 Our model experiments show that for a ’static’ biological pump, global POM cycling (i.e. POM
111 export production, rain and burial rates within the sediment) is almost entirely controlled by PO4
112 availability, with a minor increase at higher temperatures attributable to temperature-dependent
113 productivity (Fig. 2a–c). Global mean photic-zone [H2S] shows little temperature sensitivity and
−1 114 remains below 4µmol kg (Fig. 2d) and the depth of maximum [H2S] is generally below 400m
115 (Extended Data Fig. 7). At the same time, the extent of seafloor anoxia increases to values above
116 30% even at moderately high warming and nutrient increases (Fig. 2e). The relative lack of
117 sensitivity of shallow euxinia vs. deep anoxia cannot easily be reconciled with observations for
118 photic-zone euxinia (Extended Data Fig. 7) and expanded seafloor anoxia (Fig. 1b+c), respec-
119 tively.
120 In contrast, a temperature-dependent biological pump results in more efficient POM reminer-
121 alization and nutrient recycling in the upper ocean, which significantly enhances export production
122 (2f, and Extended Data Fig. 5a–d). This positive feedback is further intensified by the assumed
6 23 123 slower-than-modern sinking rate in the Paleozoic (Extended Data Fig. 5c), as it increases the
124 residence time of POM in the upper ocean, leading to more complete remineralization and nutri-
125 ent recycling. Although POM export production increases by up to a factor of three compared to
126 the static model, the POM rain to the sediments does not change significantly (Fig. 2g) due to
127 the temperature driven, efficient POM remineralization in the water column. Mean photic-zone
−1 128 [H2S] is substantially higher (reaching values > 60µmol kg ) and highly sensitive to tempera-
129 ture variations (Fig. 2i). Notably, higher temperatures cause an increase in upper ocean [H2S]
130 and a concomitant shallowing of the chemocline – even without the need to invoke any change in
131 nutrient inventory (Extended Data Fig. 8). In addition, increased ocean euxinia leads to higher
132 sulfurization rates causing POM burial rates to increase (Fig. 2h) and the extent of seafloor anoxia
133 to decrease (compare Fig. 2j+e; S6), a negative feedback that may be important in the transition
33 134 back to a more oxygenated ocean .
135 Our model predictions for the spatial distribution of euxinia and seafloor anoxia appear in
136 general agreement with the available evidence for local redox-conditions (Fig. 3) and our U-model
137 results (Fig. 1b+c). In our “Late Permian background” scenario (i.e. lower temperatures and
138 modern ocean [PO4]), only a few very local environments developed [H2S] around 20–30µmol/kg
139 and are limited to depths below 284m (Fig. 3a+d), similar to profiles in the modern Cariaco
34 140 Basin . Instances of upper ocean euxinia were very limited in spatial extent and only occured in
141 the Eastern Equatorial Panthalassic and the Eastern Tethys (South China). The simulated extent
35 142 of bottom water anoxia, 0.7% of the seafloor, is comparable to the modern value of 0.6% ( , Fig.
143 3g). Consistent with proxy reconstructions, climate warming together with a moderate release
7 144 of nutrients (i.e. the “Start Warming” phase), resulted in increasing [H2S], with sulphidic waters
145 impinging onto continental shelves and slopes, especially in warm equatorial waters (Fig. 3b+e).
146 While most of the global seafloor remains oxic (fanox =1.7%), approximately 20% of the seafloor
147 shallower than 1000m exhibits [O2] <60µmol/kg (Fig. 3h, Extended Data Table 4), a typical
148 threshold for hypoxic conditions and considered critical for the survival of many modern marine
22 149 animals . A second increase in temperature together with another slight relaxation of nutrient-
◦ 150 limitation (i.e. a ∼7 C warming of SSTs and a doubling of ocean [PO4] compared to modern) is
151 sufficient to simulate the “Main Extinction”. This scenario satisfies not only local evidence for an
152 expansion of upper ocean euxinia (Fig. 3c) but also predicts more widespread seafloor anoxia that
238 153 has spread to the abyssal plain (Fig. 3i), in broad agreement with inferences from δ U data (Fig.
154 1b+c). The habitable area in the upper ocean is further restricted as now ∼30% of the seafloor
155 above 1000m exhibits hypoxic conditions (Extended Data Table 4).
156 Implications for ocean-redox and carbon cycle dynamics
6 157 Given the relatively prolonged interval of warming in the lead up to the P/Tr boundary (∼39 kyr ),
158 our analysis favors a conceptual model in which higher metabolic rates drive a pronounced verti-
159 cal partitioning in ocean redox. This biogeochemically driven partitioning occurs within an ocean
13, 36 160 that remains not only well ventilated at depth, but consistent with previous modelling work ,
161 increases in overturning strength for warmer states (Figure 3j-l). The temperature impact on
162 metabolic rates alone is hence sufficient to provide an effective kill mechanism for the upper
163 ocean ecosystem. This is in contrast with a physical partitioning such as simulated by Penn et
12 164 al. using the Community Earth System Model (CESM), whereby instantaneous warming drives
8 165 persistent (multi kyr) stratification of the ocean. These two conceptual (and numerical) models
166 also fundamentally differ in the consequent predictions of biological export – vigorous circulation,
167 in conjunction with temperature-driven nutrient recycling substantially enhances export, whereas
168 stratification severely restricts upwelling and hence export – with important implications for the
169 stressors associated with the marine extinction. Given that the initial ocean circulation response of
170 the cGENIE ESM to instantaneous warming is very similar to that of CESM immediately after the
171 perturbation (Extended Data Fig. 10), it is clear how critical the time-scale of warming is to fully
172 understand the mechanistic link between anoxia and the extinction, highlighting the importance of
173 further work in refining age models and temperature reconstructions for the latest Permian.
174 Our inference of a metabolically-driven – rather than transient circulation-driven redox par-
4 175 titioning, allows for persistently high temperatures during the Early Triassic to prolong the oc-
9 37 176 currence of extensive seafloor anoxia . Episodic volcanism in the Early Triassic and consequent
177 oscillations between sulfidic and oxic conditions potentially played a role in the protracted biogeo-
178 chemical and biological recovery prior to the Middle Triassic. Our finding that the temperature
179 response of nutrient recycling and ocean redox is somewhat amplified in the model under a slower
180 POM sinking rate also implies that the severity of marine extinction should decrease through the
181 Phanerozoic. Specifically, the advent of pelagic biomineralization in the early Mesozoic Marine
38 23 182 Revolution , leading to a step-increase in POM sinking rates , may have decreased the sensitivity
183 of ocean anoxia to fluctuations in global climate and hence volcanism-related carbon release in the
39 184 latter half of the Phanerozoic .
9 13 185 Finally, our numerical model analysis sheds new light on the δ C decline itself, and specif-
186 ically the cause for substantive variability in what is nominally a globally imprinted signal. Dif-
187 ferences in onset timing, temporal evolution, and maximum isotopic change of the end-Permian
13 10 188 δ C decline have been observed among stratigraphic sections (e.g. the amplitude varies be-
189 tween 4 and 7h), complicating estimates of the size and source of the carbon cycle perturbation.
13 190 We simulate a maximum decrease in the Tethys Ocean subsurface δ C during the P/Tr of 3.9h
13 191 – approximately 60% of the observed δ C decrease in Armenia and Iran, with a smaller shift
192 occurring in colder waters (Fig. 1e). This decline is driven only by temperature-induced changes
193 to the biological pump in the model without invoking changes in the ocean+atmosphere carbon
194 inventory (see SI). The implication is that end-Permian carbon release as inferred from observed
13 3 195 δ C changes (e.g. Erwin ) is likely overestimated, with the total release being smaller and/or the
196 isotopic composition of the source is less negative, than previously assumed. This implies that
197 carbon from volcanism rather than a reduced source (e.g. organic matter) was more dominant. In
198 sum, these results reassess the relationship between temperature and the ocean redox state, and
13 199 carbon cycle perturbations recorded by δ C.
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16 18 δ Oapatite (‰ VSMOV)● ●
22 21 20 19● 18 ●
● ● Armenia ● Iran ● (a) ● (b) (c) SLM cubic fit ● ● LOESS fit ● ● ● ●●● U/Pb dates ●● ● ● Bed 28 ● ● ●● ● ● ● PTB ● ● Dajiang ● ● ● Zuodeng ● ● ● ● Guandao Bed 25 ● ● ● Daxiakou ● Zal ● ● Taşkent Age (Ma) ● ● Jesmond ● ● ● Induan ● Kamura ●● ● ●● ● ●●● ● ●● ● ●● ●●●●● ●● ●●● 252 251●● ● 250 ● ●●●● ●● ● ●● ● ● ●● (d) (e) Chang. Olenekian Permian Triassic Bed 22
20 25 30 35 40 −0.6 −0.2 0 10 20 30 40 252.120 252.0 25 30 251.9 35 251.8 40 -2 0 2 4 238U 13 Temperature (°C) δ fanox Temperature (°C) δ Ccarb
Figure 1: Temporal relationship of changes in isotope records and U mass balance modeling
for the P/Tr extinction. (a) Temperature reconstructions from the δ18O of biogenic apatite from
South China (black circles)4. (b) Carbonate δ238U data (colored symbols)9,17, 40–42. U isotope for- ward model results are shown, with gray lines representing individual iterations from the Monte
Carlo routine with variable isotopic fractionation into anoxic sediments, riverine δ238U, and car- bonate diagenetic offset. The black line represents the moving average of all model runs (SI). (c)
Varying perturbations in seafloor anoxic fraction (fanox) that drive the U-model in panel b (indi- vidual iterations, grey lines; average, black line). (d + e) Conodont apatite δ18O and δ13C6,43 (SI).
Grey shading in (e) represents the range of cGENIE δ13C values simulated for the Tethys Ocean for the experiments shown in Fig. 3.
17 Export production POM rain POM burial Mean photic-zone [H2S] Bottom water anoxia (PgC yr-1) (PgC yr-1) (PgC yr-1) (μmol kg-1) (%) 2.5 2.5 2.5 2.5 2.5 20 1.6 1.8 0.5 0.7 70 2 3 80 60 (a) (b) (c) 0.4 0.6 (d) (e) 1.6 1 70 50 3 1.4 3 3 3 2 60 3 2 15 2 2 0.3 0.5 2 2
modern 1 40 80 15 1.4 0.4 50 1.2 0.2 70 4 1 30 40 0.3 60 2 2 1.2 2 2 20 30 2 1.5 1.5 1 1.5 0.1 1.5 1.5 50 1 0.2 10 20 10 40 10 30
Static ModelStatic 10 0.1 Ocean PO 0.8 1 1 0.8 1 1 1 1 1 1 1 1 24 26 28 30 32 34 24 26 28 30 32 34 24 26 28 30 32 34 24 26 28 30 32 34 24 26 28 30 32 34
2.5 2.5 2.5 0.9 2.5 2.5 1.8 1 30 70 50 1.6 0.8 (f) 25 35 (g) (h) 0.9 (i) (j)30 20 60 1.6 0.7 60 20 1.4 0.8 30 50 45 0.6 0.7 40 30 3 1.43 3 3 10 3 2 2 2 0.5 0.6 2 50 2 1.2 10 20 30 modern 40 1 10 20 1.2 0.4 0.5 40 10
4 20 1 0.3 0.4 25 1 35 2 2 2 0.3 2 2 1.5 15 1.5 0.8 1.5 0.2 1.5 1.5 30 0.8 30 0.1 0.6 0.2 10 20 Ocean PO 20 1 25 1 0.6 1 0.1 1 1 Dynamic Model Dynamic 1 1 1 1 1 24 26 28 30 32 34 24 26 28 30 32 34 24 26 28 30 32 34 24 26 28 30 32 34 24 26 28 30 32 34 Tethys SST (°C) Tethys SST (°C) Tethys SST (°C) Tethys SST (°C) Tethys SST (°C)
Figure 2: Global sensitivity of particulate organic matter (POM) cycling and ocean redox to temperature and nutrient changes for the static and dynamic representation of the biological pump: POM flux values (in PgC yr−1) of export production (a+f), POM rain to the seafloor (b+g)
−1 and POM burial in the sediments (c+h). Mean photic-zone [H2S] (in µmol kg , d+i) and bottom
−1 water anoxia (e+j) as fraction of seafloor area with [O2] ≤ 5µmol kg (in %). The red diamonds indicate results for the experiments presented in Fig. 3.
18 Late Permian background Start Warming Main Extinction 25.6.°C SST Tethys 29.5°C SST Tethys 33.0°C SST Tethys
2.5 x pCO2, 1.0 x [PO4] 7.5 x pCO2, 1.5 x [PO4] 20.0 x CO2, 2.0 x [PO4] 90 90 90 60 60 60 200 (a) (b) (c) 30 30 30
200 100 100 100 200 200 100 0 0 100 200 0 100 mol kg-1) 100 200 100 200 200
S < 1000m 100 μ 2 S (
-30 -30 -30 2 H -60 -60 -60
Max H 0 -90 -90 -90 -180 -90 0 90 180 -180 -90 0 90 180 -180 -90 0 90 180 90 90 90
S 60 60 60 81 2 (d) (e) (f) 30 30 30 175 284
0 0 0 411
558 -30 -30 -30 Depth (m) 729 -60 -60 -60 Depth of Max H Depth -90 -90 -90 928 -180 -90 0 90 180 -180 -90 0 90 180 -180 -90 0 90 180
90 90 200150 90 0 15010060 60 200 50 10060 5050 05
5 5 2 60 60 60 05 200 100 60 200 100150 5060 5 60 5 (g) (h) (i) 0 150 30 30 60 30 5
200 100 5 50 5 0 5 200
100 5
200 100 150100 200 200 150100 6050 5 5 150 0 0 60 50 0 50 100 60
50 100 50 mol kg-1)
5 0 150 5 μ
50
60 60 5 5 60 50 0
50 ( 150 Latitude Latitude Latitude 0 6050 60
5 2 200 60 5 -30 150 -30 -30 6050100 200 200 100 150 50 100 150 50 O 0.7% 1.7% 60 50 29.3% Bottom water O water Bottom -60 -60 150 -60 60 0 -90 -90 150 -90 100 -180 -90 0 90 180 -180 -90 0 90 180 -180 -90 0 90 180 Longitude Longitude Longitude 0 0 0 (j) (k) (l) 15 1 1 1 10 5 2 2 2 0 5 MOC (Sv)
Depth (km) Depth 3 3 3 10
Ocean circulation 15 4 4 4 90 60 30 0 30 60 90 90 60 30 0 30 60 90 90 60 30 0 30 60 90 Latitude Latitude Latitude
Figure 3: Ocean redox conditions during the P/Tr transition using the dynamic cGENIE model: (a-c): Simulated maximum [H2S] between 81 and 928m. (d-f): Depth where the maximum in [H2S] is observed. (g-i): Simulated extent of seafloor anoxia (values indicate fanox). Model results for [H2S] and [O2] are superimposed by proxy observations: Evidence for euxinia/anoxia is represented by circles; evidence against by crosses; ambiguous evidence or dynamic redox- conditions by triangles (see Extended Data Table 2). (j-l): Meriodional overturing circulation
(MOC, in Sv).
19 306 Acknowledgements We gratefully acknowledge data contributions from M. Joachimski and advice on
307 how to align proxy-records. D.H. is supported by a postdoctoral fellowship from the Simons Foundation.
308 S.J.V. is a NASA Postdoctoral Program fellow. D.H. and A.R. acknowledge support from the Heising-
309 Simons Foundation.
310 Author contributions D.H., K.V.L. and A.R. conceived the study; D.H. and A.R. designed and conducted
311 cGENIE experiments; K.V.L. adapted the U-model and conducted the experiments; K.V.L. and D.H. com-
312 piled and analyzed the proxy-data; All authors analyzed model output. D.H., K.V.L. and A.R. wrote the
313 manuscript with input from all authors.
314 Competing Interests The authors declare no competing interests.
315 Supplementary Information is available for this paper.
316 Correspondence and requests for materials should be addressed to D.H. (email: [email protected]).
317 Data availability
318 The locations of all data used in this study are provided in the supplementary information.
319 Code availability
320 The version of the code used in this paper is tagged as release v0.9.15 and has a DOI of 10.5281/zen-
321 odo.4008865. Necessary boundary condition files are included as part of the code release. Con-
322 figuration files for the specific experiments presented in the paper can be found in the installation
323 subdirectory: genie-userconfigs/MS/huelseetal.2020. Details of the experiments, plus the com-
324 mand line needed to run each one, are given in the readme.txt file in that directory. A man-
20 325 ual describing code installation, basic model configuration, and an extensive series of tutorials
326 is provided. The Latex source of the manual and pre-built PDF file can be obtained by cloning
327 (https://github.com/derpycode/muffindoc).
21 Figures
Figure 1
Temporal relationship of changes in isotope records and U mass balance modeling for the P/Tr extinction.
Figure 2
Global sensitivity of particulate organic matter (POM) cycling and ocean redox to temperature and nutrient changes for the static and dynamic representation of the biological pump. Figure 3
Ocean redox conditions during the P/Tr transition using the dynamic cGENIE model.
Supplementary Files
This is a list of supplementary les associated with this preprint. Click to download.
SIHuelseeaEndPermian.pdf